STAR 1

on

THE TECTONIC EVOLUTION

OF GREENSTONE BELTS

(NASA-Cfi-176619) HORKSHOP ON "THE TECTONIC

EVOLUTION OF' GEEENSTONE BEI1S (lunar and

Planetary Inst,) 143 p HC AC7/MI A01

CSC1 08G

W86-23089

THRU

M86-23137

Unclas

G3/46 057.17

HOUSTON, TEXAS

JANUARY 16-18,1986

A Lunar and Planetary Institute Workshop

UNIVERSITIES SPACE RESEARCH ASSOCIATION

LUNAR AND PLANETARY INSTITUTE

3303 NASA ROAD 1

HOUSTON, TEXAS 77058
WORKSHOP ON THE TECTONIC EVOLUTION OF GREENSTONE BELTS

HOUSTON, TEXAS

January 16 - 18, 1986

A Lunar and Planetary Institute Workshop

Compiled by:

Publications Office

Lunar and Planetary Institute

3303 NASA Road One

Houston, TX 77058-4399

LPI Contribution 584
Material in this volume may be copied without restraint for abstract

services, libraries, and individual research purposes. Republication

of any abstract or portion thereof requires written permission from

the author as well as appropriate acknowledgement of this publication,

********* *
CONTENTS

Hot Spot Abundance, Ridge Subduction and The Evolution of

Greenstone Belts

D. Abbott and S. Hoffman 1

Lithology, Age and Structure of Early Proterozoic Greenstone

Belts, West African Shield

K. Attoh 4

Correlations and Contrasts in Structural History and Style

Between an Archean Greenstone Belt and Adjacent Gneiss Belt,

NE Minnesota

R. L. Bauer, P. J. Hudleston, and D. L. Southwick 7

Tectonic Evolution of Greenstone-Gneiss Association in Dharwar

Craton, South India: Problems and Perspectives for Future

Research

Y. J. Bhaskar Rao 9

Can Trace Element Distributions Reclaim Tectonomagmatic Facies

of Basalts in Greenstone Assemblages?

J. C. Butler 11

Barberton Greenstone Belt Volcanism: Succession, Style, and

Petrogenesis

G. R. Byerly and D. R. Lowe 14

Tectonic Setting and Evolution of Late Archean Greenstone Belts

of Superior Province, Canada

K. D. Card 17

High Precision U-Pb Geochronology and Implications for the

Tectonic Evolution of the Superior Province

D. W. Davis, F. Corfu, and T. E. Krogh 20

Crustal Structure of the Archean Granite-Greenstone Terrane in

the Northern Portion of the Kaapvaal Craton

J. H. de Beer, E. H. Stettler, J. M. Barton, Jr., D. D. van Reenen

and J. R. Vearncombe 23

Extensional Tectonics During the Igneous Emplacement of the

Mafic-Ultramafic Rocks of the Barberton Greenstone Belt

M. J. de Wit 25

A Mid-Archean Ophiolite Complex, Barberton Mountain Land

M. J. de Wit, R. Hart, and R. Hart 27

Felsic Igneous Rocks Within the Barberton Greenstone Belt:

High Crustal Level Equivalents of the Surrounding TonaliteTrondhjemite Terrain Emplaced During Thrusting

M. J. de Wit and A. H. Wilson 30

Heat Flow and Heat Generation in Greenstone Belts

M. J. Drury 33
Komatiite Genesis in the Archean Mantle, with Implications for

the Tectonics of Archean Greenstone Belts

D. Elthon 36

The Yilgarn Craton Western Australia: A Tectonic Synthesis: I

R. E. P. Fripp 39

Tectonics of Some Amazonian Greenstone Belts

A. K. Gibbs 41

Spatial Greenstone-Gneiss Relationships: Evidence from

Mafic-111 tramafic Xenolith Distribution Patterns

A. Y. Glikson 44

Dismembered Archean Ophiolite in the Southeastern Wind River

Mountains, Wyoming — Remains of Archean Oceanic Crust

G. D. Harper 47

Preliminary Report on the Geology and Gold Mineralization of

the South Pass Granite-Greenstone Terrain, Wind River

Mountains, Western Wyoming (USA)

W. D. Hausel 50

The Kolar Schist Belt: A Possible Archean Suture Zone

G. N. Hanson, E. J. Krogstad, V. Rajamani, and S. Balakrishnan 52

Evidence for Spreading in the Lower Kam Group of the Yellowknife

Greenstone Belt: Implications for Archean Basin Evolution in the

Slave Province

H. Helmstaedt and W. A. Padgham 55

Archaean Wrench-Fault Tectonics in the Abitibi Greenstone Belt

of Canada

C. Hubert and J. N. Ludden 59

Transpression as the Main Deformational Event in an Archean

Greenstone Belt, Northeastern Minnesota

P. J. Hudleston, D. Schultz-Ela, R. L. Bauer, and D. L. Southwick 62

A Continuous Record of Tectonic Evolution from 3.5 Ga to 2.6 Ga

in Swaziland and Northern Natal

D. R. Hunter, A. H. Wilson, J. A. Versfeld, A. R. Allen,

R. G. Smith,D. W. W. Sleigh, P. B. Groenewald, G. M. Chutter,

and V. A. Preston 64

Is the Cameron River Greenstone Belt Allochthonous?

T. M. Kusky 69

Sedimentological and Stratigraphic Evolution of the Southern

Part of the Barberton Greenstone Belt: A Case of Changing

Provenance and Stability

D. R. Lowe and G. R. Byerly 72

Evidence for Structural Stacking and Repetition in the

Greenstones of the Kalgoorlie District, Western Australia

J. E. Martyn 75
Preliminary Structural Model for the Southwestern Part of

the Michipicoten Greenstone Belt, Ontario

G. E. McGill and C. H. Shrady 78

Thermal Implications of Metamorphism in Greenstone Belts and

the Hot Asthenosphere-Thick Continental Lithosphere Paradox

P. Morgan 81

Geochemical Characters and Tectonic Evolution of the Chitradurga

Schist Belt: An Archaean Suture(?) of the Dharwar Craton, India

S. M. Naqvi 84

Geochemical and Isotopic Constraints on the Tectonic Setting of

the Serra dos Carajas Belt, Eastern Para, Brazil

W. J. Olszewski, Jr., A. K. Gibbs, and K. R. Wirth 86

Polyphase Thrust Tectonics in the Barberton Greenstone Belt

I. Paris 89

Age Constraints on the Evolution of the Quetico Belt, Superior

Province, Ontario

J. A. Percival and R. W. Sullivan 91

Archean Megacrystic Plagioclase Units and the Tectonic Setting

of Greenstones

W. C. Phinney, D. A. Morrison, and D. Maczuga 94

Rainy Lake Wrench Zone: An Example of an Archean Subprovince

Boundary in Northwestern Ontario

K. H. Poulsen 97

A Palaeomagnetic Perspective of Precambrian Tectonic Styles

P. W. Schmidt and B. J. J. Embleton 100

The Wisconsin Magmatic Terrane: An Early Proterozoic

Greenstone-Granite Terrane Formed by Plate Tectonic

Processes

K. J. Schulz and G. L. LaBerge 102

New Insights into Typical Archean Structures in Greenstone

Terranes of Western Ontario

W. M. Schwerdtner 105

Deformational Sequence of a Portion of the Michipicoten

Greenstone Belt, Chabanel Township, Ontario

C. H. Shrady and G. E. McGill 106

A Continental Rift Model for the La Grande Greenstone Belt

T. Skulski, A. Hynes, M. Liu, D. Francis, B. Rivard, and

K. Stamatelopoulou-Seymour 108

Two Contrasting Metamorphosed Ultramafic-Mafic Complexes

from Greenstone Belts, the Northern Kaapvaal Craton and

Their Significance in Archaean Tectonics

C. A. Smit and J. R. Vearncombe 111
Zircon Lu-Hf Systematics: Evidence for the Episodic Development

of Archean Greenstone Belts

P. E. Smith, M. Tatsumoto, and R. M. Farquhar 114

Early Precambrian Crustal Evolution of South India

R. Srinivasan 117

Basement-Cover Relations and Internal Structure of the Cape

Smith Klippe: A 1.9 Ga Greenstone Belt in Northern Quebec,

Canada

M. R. St-Onge, P. F. Hoffman, S. B. Lucas, D. J. Scott, and

N. J. Begin 118

Rhyolitic Components of the Michipicoten Greenstone Belt,

Ontario: Evidence for Late Archean Intracontinental Rifts

or Convergent Plate Margins in the Canadian Shield

P. J. Sylvester, K. Attoh and K. J. Schulz

Geophysical Characteristics and Crustal Structure of

Greenstone Terranes, Canadian Shield

M. D. Thomas, L. Losier, P. C. Thurston, V. K. Gupta,

R. A. Gibb, and R. A. F. Grieve 123

Volcanological Constraints on Archean Tectonics

P. C. Thurston and L. D. Ayres 126

The Dehydration, Rehydration and Tectonic Setting of

Greenstone Belts in a Portion of the Northern Kaapvaal

Craton, South Africa

D. D. van Reenen, J. M. Barton, Jr., C. Roering,

J. C. van Schalkwyk, C. A. Smit, J. H. de Beer, and

E. H. Stettler 129

The Stratigraphy of the Steep Rock Group, N.W. Ontario,

with Evidence of a Major Unconformity

M. E. Wilks and E. G. Nisbet 132

Author Index 135
N86-23090

HOT SPOT ABUNDANCE, RIDGE SUBDUCTION AND

THE EVOLUTION OF GREENSTONE BELTS

Dallas Abbott and Sarah Hoffman

College of Oceanography

Oregon State University

Corvallis, Oregon 97331

A number of plate tectonic hypotheses have been proposed to explain the

origin of Archaean and Phanerozoic greenstone/ophiolite terranes. In these

models, ophiolites or greenstone belts represent the remnants of one or more

of the following: island arcs (1,2), rifted continental margins (3), oceanic

crustal sections (1,4), and hot spot volcanic products (1,3,5). If plate

tectonics has been active since the creation of the earth, it is logical to

suppose that the same types of tectonic processes which form present day

ophiolites also formed Archaean greenstone belts. However, the relative importance of the various tectonic processes may well have been different.

The Archaean earth is postulated to have had greater internal heat production and consequently a younger maximum age of the oceanic lithosphere at

subduction (6,7). One of the consequences of a greater proportion of subduction of young oceanic lithosphere in the Archaean is that ridge subduction

would have been more common (7). The most common type of ridge subduction

in the Archaean would have been that where oceanic lithosphere comprised both

the overriding and subducting plate. The only present day example of this

type of subduction is the subducting ridge in the Woodlark basin. This ridge

crest has several geochemical anomalies: basalts with an island arc signature,

and a dacite volcano on the ridge crest (8,9). The island arc component of

the basalts has two proposed origins: contamination by an older subducting

plate due to polarity reversal of the arc (9) and fluid contamination from

the base of the subducting plate (10). Plate reorganization and ridge subduction are both postulated to have been more abundant in the Archaean (7).

Regardless of the mechanism by which the arc-like component is generated,

Archaean oceanic crust emplaced on land would have been much more likely to

have an arc-like composition. Similarly, the dacite volcano observed on the

Woodlark basin ridge crest could also have counterparts in Archaean greenstone belts.

Other aspects of the Woodlark basin subduction system may also have relevance for Archaean greenstone belts. The New Georgia island arc, which is

being formed by subduction of the oceanic crust of the Woodlark basin (Figure

I)/ is composed of overlapping volcanoes, located 4-70 km above the Benioff

zone (11,12). The New Georgia arc is quite different from a 'typical1

Phanerozoic arc, e.g. the Marianas arc (Figure 2). In the Marianas, the

volcanoes are spaced 50-100 km apart and sit 125-150 km above the Benioff

zone (13,14). The island arc volcanics of the New Georgia arc also have some

unusual characteristics. One island is a picritic volcano, thought to be

the direct result .df the ridge subduction process (8) . If a higher percentage of Archaean island arcs were like the New Georgia islands, individual

volcanoes would possess overlapping edifices and picritic volcanoes would

occasionally occur. The overlapping volcanic edifices would increase the

thickness of layer 2 (the pillow basalt layer) and would increase the probability of multiple phases of hydrothermal activity. Consequently, the relative abundance of Archaean ore deposits could be due to the greater incidence of New Georgia-like island arcs.
HOT SPOTS AND RIDGE SUBDUCTION

2 Abbott, D. and Hoffman, S.

Another probable consequence of greater internal heat production in the

Archaean would have been a greater abundance of hot spot activity. For example, in the Phanerozoic, global ridge volume in the Cretaceous is thought

to have increased and to have caused the Cretaceous sea-level high. This increase in the global sea floor creation rate may have coincided with an increase in hot spot activity (15). If increases in hot spot activity do

coincide with increases in sea floor creation rate, hot spot activity must

have been much more abundant in the Archaean. At present, 10% of all sea

floor volcanism is estimated to result from hot spot activity (16). In the

Archaean, it is likely that an even greater percentage of sea floor magmatism would have been hot spot generated.

Greater hot spot magmatism in the Archaean would have increased the incidence of bouyant subduction. Bouyant subduction can be a result of subduction of young oceanic crust or of older oceanic crust with a thickened

crustal section (7). Much of the oceanic crust which subducts bouyantly has

no volcanism or reduced volcanism. This reduction in volcanic activity as a

result of bouyant subduction is most common if the overlying plate has a

thickened crustal section. Consequently, an increase in hot spot activity

in the Archaean could have decreased the percentage of subducting plates

causing magmatic activity in the overriding plate, particularly when the

overriding plate was relatively cold, thick continental lithosphere.

Areas of hot spot magmatism generally have a thickened pillow basalt

section and a greater abundance of highly permeable rocks. These thickened

pillowed sections can support more intense hydrothermal activity. Increased

hydrothermal alteration at hot spots, particularly ridge-centered hot spots,

could also have contributed to the relative abundance of Archaean massive

sulfide deposits.

In conclusion, it is probable that many of the differences in preserved

Archaean and Phanerozoic greenstone belt/ophiolite terranes can be explained

as a result of a difference in the relative importance of different plate

tectonic processes. This difference is a direct result of the increased internal heat production of the earth in the Archaean.

REFERENCES: (1) Burke, K. , J. F. Dewey, and W. S. F. Kidd (1976), The Early

History of the Earth, Windley (ed.), 113-129; (2) Miyashiro, A. (1973),

Earth Planet. Sci. Lett., 19, 218-224; (3) Clarke (1970), Contrib. Mineral.

Petrol., 25, 203-224; (4) Gass, I. G. (1977), Volcanic Processes in Ore

Genesis, Geol. Soc. of London, 72-77; (5) Schulz, K. J. (1977), University of

Minnesota, Ph.D. Thesis, 349 pp.; (6) Bickle, M. J. (1978), Earth Planet.

Sci. Lett., 40, 301-315; (7) Abbott, D. H., and S. E. Hoffman (1984),

Tectonics, 3_, 429-448; (8) Johnson et al., AAPG CircumPacific Energy Series,

in press, 1985; (9) Perfit et al., AAPG CircumPacific Energy Series, in

press, 1985; (10) Abbott, D. H., and M. Fisk, in preparation, 1985;

(11) Taylor et al. , (1982), EOS, 6^3, 1120-1121; (12) Cooper, P. A., and

B. Taylor (1985), Nature, 314, 428-420; (13) Karig, D. (1971), J. Geophys.

Res., 76, 2542-2560; (14) Karig, D. (1971), Init. Rpts. DSDP, £, Fischer,

A. G., Heezen, B. C. et al., 681-689; (15) Ribe, N. M., and A. B. Watts

(1982), Geophys. J. R. Astr. Soc., 71, 333-362; (16) Kennett, J. P. (1982),

Marine Geology, Prentice-Hall, Inc., 813 pp.
i Q n

'

HOT SPOTS AND RIDGE SUBDUCTION

Abbott, D. and Hoffman, S.

155 156 157

100-

200

300

SCT NST

100 200 300 400

KEY

A DREDGE LOCATION

TRENCH AXIS

RIDGE CREST

& 2 M.Y OLD

> 2 MY. OLD

155 156 157

Figure 1. (left) Benioff zone of the New Georgia arc (SCT), after (12);

T = Trench, V = Volcanic Line. (right) Volcanoes of the New Georgia arc,

after (10).

USE

IMC

MCIHC OCCAM

Figure 2. (left) Benioff zone of the Marianas arc, after (13); T and V as in

Figure 1. (right) Volcanoes of the Marianas arc are designated by dots,

after (14).
N86-23091

LITHOLOGY, AGE AND STRUCTURE OF EARLY PROTEROZOIC

GREENSTONE BELTS, WEST AFRICAN SHIELD; Kodjo Attoh, Geology

Department, Hope College, Holland, MI 49423

Distribution, Lithologic characteristics and Stratigraphic

relations. Distribution of early Proterozoic volcanic rocks in

the West African shield is shown in Fig. 1; an approximate

boundary between Archean age terrane, to the west, and the

Proterozoic terrane to the east, is partly marked by a major

fault. Lithologic and chemical data have been compiled for belts

(2-91 in the Proterozoic terrane from BRGM reports {_!, 2J .

Available stratigraphic information from geologic maps of these

areas indicate that a typical sequence is comprised of

predominately mafic lava flows (basalt-andesite) at the base,

which are overlain by felsic volcanic rocks including pyroclastic

rocks and lavas. This succession, referred to as Lower Birimian,

is overlain by Middle and Upper Birimian sedimentary rocks.

Lithostratigraphic data from belt (1), located in northeastern

Ghana [3] , indicate the volcanic succession is 6-8 km thick. The

lowest unit in this succession is represented by 2 km of felsic

pyroclastic rocks, flows and fine grained sediments. This is

followed by 3-4 km of basaltic lava flows which are locally

pillowed, the top of the unit is marked by a distinctive

manganese formation (MF) consisting of Mn-Fe rich cherts up to

200 m thick. Dacitic lithic tuffs, welded tuffs and andesitic

flows up to 2500 m thick overlie the mafic lava flow unit. The

youngest volcanic unit consists of mafic tuffs and breccia with a

distinctive fragmental texture. Preliminary data indicate that a

similar succession occurs in belt (10). The internal plutonic

rocks of belt (1) include: (a) hornblende-bearing granodiorite

bodies considered to be subvolcanic plutons (CT-plutons) ; and (b)

post-kineraatic mica-bearing granitic plutons (pi -plutons).

External plutonic rocks include tonalitic and granodioritic rocks

which immediately flank the volcanic belt, and paragneisses which

occur within the plutonic terrane.

Chemical charcteristics and Ages. Of about 100 chemical

analyses reported for belts(2-9) calc-alkaline rocks consitute

55% and tholeiites 45%. Quartz-normative basalt constitutes 99%

of the rock type in the tholeiitic suite. In the calc-alkaline

suite, 9% of the analyses is basalt, 45% andesite and the rest is

dacite and rhyodacite. The ratio of tholeiitic to calc-alkaline

rocks based on the stratigraphic thicknesses and chemical

analyses of samples from belt (1) is between 57% and 43%.

Ultramafic volcanic rocks occur in belt (3), indicated from

chemical data from belt (6) and (9) and constitute 1% of all

samples analysed. Komatiites have not been reported from the

West African Shield, thus the rocks analysed may be classified as

high-Mg-basalts. The tholeiitic rocks from belt (1) are enriched

in Ti, and depleted in Zr relative to modern ocean floor basalts

[l] , and are depleted in K, Rb, Sr and Ba relative to the

calc-alkaline rocks. Within the calc-alkaline suite which

include the subvolcanic plutons, the major and trace elements

show continuous trends from calc-alkali basalts to rhyolites.
EARLY PROTEROZOIC GREENSTONE BELTS, WEST AFRICA

Attoh, K. '

The hornblende-bearing plutons plot in granodiorite and diorite

fields of Q-Kf-Pl diagram; whereas the rocks from the pi-plutons

have normative and modal mineral compositions of granodiorite,

quartz-monzonite and minor quartz synite and monzonite. All the

plutonic rocks are strongly HREE depleted \S]. The -plutons

(Si02=56-66*) show the least depletion with [La/Yb]n = 13-43. The

paragneisses of the external plutonic terrane (Si02 = 70-7l3>) show

the steepest REE pattern with [La/Yb]n = 33 - 66; while the

post-kinematic plutonic rocks (Si02 = 70-75), and La/Yb = 18 - 58,

are somewhere in between. Relative to the subvolcanic plutons

with (Th=1.9-5.7, and U=0.9 -1.9) the pi-plutons are enriched in

Th and U (Th=7.7-10.9 and U-4.5-25ppm). Age of volcanism in the

West African Shield is not known; however, K/Ar and Rb/Sr ages

have been reported for the rocks which intrude the volcanic rocks

and can be used to place minimum age limits. Rb/Sr analyses of

mica pi-pluton samples from belts (2-9) yielded the following

ages (my): 1870+157 to 2004+42 for whole rock; and 1940145

mineral (plagioclase) isochron [5} . K/Ar analyses of amphiboles

from belt (1) gave the following ages: (i) an older age of 22231

283 was obtained from hornblende in the youngest volcanic unit;

and (ii) a younger age, 2087+ 138 was obtained from zoned,

titaniferous hornblende in a deformed diorite porphyry intruded

into the lowest unit in the volcanic succession. The available

data lead to the conclusion that the minimum age for the volcanic

activity must be between 2200 and 2100 my. It is significant

that Archean ages have not been reported from any of the volcanic

belts (1-10).

Structure of an early Proterozoic Volcanic belt in

northeastern Ghana. Cleavage in the volcanic belt strikes N40E

and dips steeply to the NW and SE. Mesoscopic folds, with

locally well developed axial surface cleavage parallel to this

foliation, plunge steeply NW and SE. Because the orientation of

fold axes and cleavage surfaces do not change with respect to the

stratigraphic position, it is concluded that the whole volcanic

succession was deformed during a pre-2000 my old orogenic event.

Evidence for multiple deformation occurs in the form of NW

plunging folds and the folded trace of the axial surface of the

major folds. The strong NE-SW orientation of the major

structures is such that one has to conclude that the second

deformation was not as intense as the first. Foliation in the

external plutonic terrane is subparallel to the foliation in

adjacent volcanic rocks. Unequivocal evidence for pre-greenstone

belt structure was not found in the external plutonic terrane;

however, NS structures occur in the paragneisses, which are

oblique to NNE-NE structures in the volcanic belt. Gravity

anomalies associated with the greenstone belt and the internal

plutons have been modelled taking the surrounding plutonic

terrane as background. The model predicts that the depth to the

bottom of the volcanic succession is 3-4 km. Fig 2 is a

structural section of belt(l) based on gravity models especially

with regard to allowable geometries of the rock units at depth.

The overturned limb of the major anticlinal fold is consistent

with available facing indicators.
ORIGINAL PAGE IS

OF POOR QUALfTY

EARLY PROTEROZOIC GREENSTONE BELTS, WEST AFRICA

Attoh, K.

Fig 1. Meat African Shield showing the . distribution of

Proterozoic volcanic-sedimentary belts: 1) early

Proterozoic volcanic belts, numerical labels referred to

in text; 2) late Proterozoic platform sediments; 3)

boundary between Archean and Proterocoic shields.

Fig 2. Geologic section across belt (1) in northeastern Ghana:

1) epiclastic sediments and tuffs; 2) mafic lavas

(tholeiitic basalts); 3) felsic tuffs and intermediate

lavas (calcalkaline); 4) postkinematic granites

(pi-granite); 5) granodiorites, tonalites and

paragneisses of external plutonic terrene.

References - [l] BRQN: Bureau de Recherches Geologique et

Minieres. [2l Alsac, C., (1968, 1969) - BRGN publication G4S14;

Blanchot, A. et al (1972) Carte Geologique de la partie

meridionale de 1'Afrigue de 1'ouest (BRGN map); Doucelier, J.,

(1963) BRGM men 10; Papon, A., (1973) BRGH mem 80. [3] Attoh,

K., (1980), Geol. Surv. Can. 80-1C; (1982) Prec. Res. 18,

275-290. [O Pearce, J. A. and Cann, J. R., (1973) - Earth Plan.

Sci. Lett., 19, 290-300. L&1 Beasoles, G., (1977) BRGM mem 88.

[s] REE analyses done by Schultz, K. J.
N86-23092

CORRELATIONS AND CONTRASTS IN STRUCTURAL HISTORY AND

STYLE BETWEEN AN ARCHEAN GREENSTONE BELT AND ADJACENT GNEISS

BELT, NE MINNESOTA; R.L. Bauer, Dept. of Geology, University of

Missouri, Columbia, MO 65211, P.J. Hudleston, Dept. of Geology

and Geophysics, University of Minnesota, Minneapolis, MN 55455,

and D.L. Southwick, Minnesota Geological Survey, St. Paul, MN

55114.

Introduction. Low-grade metagraywacke and greenstone of the

Vermilion district and amphibolite facies schist and migmatites

of the Vermilion Granitic complex (VGC) are separated by a series

of east-trending dip-slip and strike-slip faults (1). Structural

analysis in the boundary region between these two terranes

indicates that they both sustained an early D.< deformation which

lead to recumbent folding. This was followed by a north-south

transpression that resulted in the generation of upright F£ folds

and locally well-developed, dextral, D2 shear zones (2). Despite

these correlations, there are distinct differences in structural

style and late-stage fold history between the two terranes that

we attribute to: 1. differences in the crustal levels of the

two terranes during deformation, and 2. effects of late-D«

plutonism in the VGC.

Do deformation produced a series of upright F~ folds with

easterly striking axial planes that are the most prominent fold

structures in both terranes. The largest fold of this series is

a westerly plunging antiform that straddles the dip-slip fault

boundary between the two terranes. Large-scale parasitic folds

on this structure are invariably of S symmetry in the southern

VGC and occur on the northern limb of the antiform. D2 dextral

shear zones are well represented in the Vermilion district where

they are generally parallel to the regional F~ axial planes.

Although distinct ductile shear zones are not observed in the

VGC, evidence of a D2 dextral shear component is locally

indicated by asymmetrical pull-aparts and rotated vein segments

in the migmatites.

F^ recumbent folding is inferred from structural facing in

the ma~jor F2 antiform that crosses the boundary between the two

terranes. Facing is downward on both limbs of the fold which is

interpreted to be part of the lower, overturned limb of a largescale F., recumbent fold. A change to upward facing strata

further south in the Vermilion district indicates a crossing onto

the upper limb of this structure. Finite strain data, determined

from clasts in sedimentary/volcaniclastic units in the Vermilion

district, can be completely accounted for in terms of the

deformation producing the F^ folds (3). Locally intense F.

folding in these rocks is therefore attributed to deformation in

soft or very poorly lithified sediment. However, biotite schists

making up part of the same structure in the VGC display a

pronounced S-, foliation that developed parallel to bedding during

the early stages of metamorphism. We have suggested that

metamorphic dehydration reactions occurring in the lower strata

led to the development of high pore pressures in the upper

portion of the sedimentary pile (4 and 5). The combination of

high pore pressures and gravitational instability during the F1
CORRELATIONS AND CONTRASTS

Bauer, et al.

folding resulted in soft-sediment, coherent down-slope movement

in the upper strata while the lower strata underwent strain and

metamorphic rescrystallization during F., folding. Soft-sediment

F., folding in the Vermilion district could have led to a rather

complex distribution of F^ structures, because the more competent

greenstones could not have been soft and therefore may have

undergone a much different response to the F., folding.

Fg folding has been observed only in the VGC to the north of

the boundary zone with the Vermilion district, near the

southwestern contact between the migmatites and the Lac La Croix

Granite batholith. Along this margin of the batholith, F2 folds

were reoriented during the emplacement of the pluton and

subsequently refolded by F<> conical folds that formed during the

waning stages of the regional north-south transpression that

generated the Y~ folds. Such F^ folds are not observed along the

southern margin of the batholitn where the F~ folds are parallel

to the batholith boundary and therefore were not reoriented.

In summary, our analysis of the deformation along the

boundary between the Vermilion Granitic Complex and the Vermilion

district indicates that the two terranes have seen a similar

deformation history since the earliest stages of folding in the

area. Despite this common history, variations in structural

style occur between the two terranes, such as the relative

development of D., fabrics and D2 shear zones, and these can be

attributed to differences in the crustal levels of the two

terranes during the deformation. Similarly, the local

development of F., folds in the VGC, but not in the Vermilion

district, is interpreted to be a result of late-Dj pluton

emplacement which was not significant at the level of exposure of

the Vermilion district.

References:

1. Sims, P.K. (1976) Bull. Geol. Soc. Amer., 87, p. 49-62.

2. Hudleston, P.J., Schultz-Ela, D., Bauer, R.L., and Southwick,

D.L., (1986) In Workshop on the Tectonic Evolution of

Greenstone BeTTs"Lunar and Planetary Institute, Houston,

(this volume).

3. Hudleston, P.J. (1976) Can, J_._ Earth Sci., 13, p. 587-592.

4. Bauer, R.L. and Hudleston, P.J. (1982) Trans. Am. Geophy.

Union, 63, p. 614.

5. Bauer, R.L. (1985) Geology, 13, p. 657-660.
N86-2309 3

TECTONIC EVOLUTION OF GREENSTONE-GNEISS ASSOCIATION IN

DHARWAR CRATON, SOUTH INDIA: PROBLEMS AND PERSPECTIVES

FOR FUTURE RESEARCH

Y.J. BHASKAR RAO, National Geophysical Research Institute, Hyderabad, INDIA

The two fold stratigraphic subdivision of the Archean-Proterozoic greenstonegneiss association of Dharwar craton into an older "Sargur group" (older than

2.9 Ga.) and a younger "Dharwar Supergroup" (1) serves as an apriori stratigraphic

model. The concordannt greenstone (schist)-gneiss (Peninsular gneiss) relationships,

ambiguities in stratigraphic correlations of the schist belts assigned to Sargur

group and difficulties in deciphering the older gneiss units can be best appreciated

if the Sargur group be regarded as a trimodal association of: (i) ultrabasic-mafic

metavolcanics (including komatiites), Til) clastic and nonclastic metasediments

and ^paragneisses and (iii) mainly tonalite/trondhjemite gneisses and migmatites

of diverse ages (2) which could be as old as c. 3.4 Ga. or even older. The extensive

occurrence of this greenstone-gneiss complex is evident from recent mapping

in many areas of central and southern Karnataka State.

The Dharwar Supergroup is deposited unconformably over an ensialic basement

comprising the older trimodal association and is further divisible into a lower

Bababudan and an upper Chitradurga groups. The volcanic and sedimentary rocks

in the Dharwar schist belts display highly variable compositions, lithofacies and

stratigraphic thicknesses. The available data is compatible with their deposition

in a variably subsiding and progressively evolving basin(s) in an intracratonic or

continental margin setting. The Bababudan group is dominated by sediments characteristic of the nearshore intratidal to shallow marine environments and subaerial

toshallow marine volcanics (3, 4). The sediment thickness and way-up criteria

are suggestive of progressive subsidence of the basin from south to north and

concomitant accumulation of sediments derived from both intrabasinal and exterior

sources which culminated in the deposition of thick (over 5 kms) sequence of

polymict conglomerates and alluvial fan deposits in the rapidly subsiding Kaldurga

basin (4). Subsequent sedimentation and volcanism proceeded in essentially deep

marine environment as evident from rocks in the interiors of Shimoga and Chitradurga belts. The volcanic character evolved from predominantly tholeiitic (with

minor komatiitic occurrences) in the lower units of Bababudan group to calc-alkalic

affinities in the upper units of the Dharwar Supergroup. The overall major and

trace element compositions of the Dharwar metavolcanics are comparable to

Phanerozoic volcanics from continental margin or back-arc settings. While both

light REE depleted and enriched types are noted often within the same volcanic

formation, an important feature of the metavolcanics is their high Zr/Y character

compared to most other Archean volcanic suites in the southern hemisphere suggesting possible trace element heterogeneities in the source regions of Dharwar

volcanic rocks (5, 6).

The greenstone and gneiss formations throughout the craton show evidences

of two or three phases of deformation with superposed folding resulting often

in complex interference patterns. Both pre-Dharwar and Dharwar formations

display broadly similar deformation styles and a remarkable parallelism in their

tectonic fabrics differing in the intensities of deformation and grade of regional

metamorphism (4, 7). The older sequences show superposition of tight upright

or overturned isoclinal and/or recumbent folds of the first and second generations

(Fl and F2) and a set of open folds (F3) and metamorphosed to amphibolite or

granulite facies while the Dharwar rocks are generally in greenschist facies with

large scale recumbent and tight isoclinal folds being uncommon (4). The structural

history of the craton is complicated by repeated syn or late tectonic diapirism
10

and intense shearing, strike-oblique slip movements and thrusting particularly

along several of the N-S trending regional shear systems (8).

Apart from the general problems concerning the conceptual approaches to

early Archean tectonics and crustal evolution, the stages of the tectonic evolution

in the Dharwar craton are poorly constrained by lack of information on many

crucial aspects of the geology such as; chrono-stratigraphy of schist belts, timing

of the major thermal and tectonic events, schist-gneiss relationships and their

relative antiquities in the (older) trimodal association, the nature and evolution

of the low grade-high grade transitions in the craton. Thus, while the evolution

of the pre-Dharwar greenstone-gneiss association is largely enigmatic, the Dharwar

Supergroup appears to be a consequence of wide-spread heating of the continental

crust around c. 3.0 Ga., tectonic instability resulting in rifting probably along

reactivated pre-existing lineaments, formation of broad basin(s), volcanism and

sedimentation concomitant with variable rates of subsidence of the basin(s) in

response to basement instability and differential upliftment of the surrounding

basement highs (horsts?) across the boundary faults (*). The tectonic evolution

of the pre-Dharwar crust and the relative importance of the "thick skin" vis-a-vis

"thin skin" tectonics (4, 8) to the Archean/Proterozoic history of the Dharwar

craton can be assessed only after more detailed structural data on a regional

scale become available in conjunction with precise and reliable data on the primary

and metamorphic ages of the schists and gneisses in the craton.

REFERENCES

1. Swami Nath 3. et_al_. (1976) Rec. Geol. Surv. India, 107(2), p. 1^3-175.

2. Bhaskar Rao Y.3. et al (1983) Memoir. Geol. Soc. India, 4, p. 309-328.

3. Bhaskar Rao Y.J. and Naqvi S.M. (1978) in B.F. Windley and S.M. Naqvi (Edt),

Archaean Geochemistry, Elsevier.

it. Chadwick B. et^ al. (1985) Jour. Geol. Soc. India Parts I and II, 26, p. 769-822.

5. Bhaskar Rao Y.3. and Drury S.A. (1982) Jour. Geol. Soc. India, 23, p. 1-12.

6. Drury S.A. (1983) Geochim. Cosmochim. Acta., 47, p. 317-329.

7. Naha K. et al. (1985) Tectonophysics (in press).

8. Drury S.A. e_^ al. (198») Jour. Geology., 92, p. 3-20.
N86-23094 11

CAN TRACE ELEMENT DISTRIBUTIONS RECLAIM TECTONOMAQMATIC

FACIES OF BASALTS IN GREENSTONE ASSEMBLAGES?

Butler, J.C. Qeosciences, University of Houston-University

Park, Houston, Texas 77004

During the past two decades many words have been written

both for and against the hypothesis that the tectonic setting of

a suite of igneous rocks is retained by the chemical variability

within the suite. For example, Pearce and Cann (1) argued that

diagrams can be constructed from modern/recent basalt

subcompositions within the system Ti-Zr-Y-Nb-Sr such that

tectonomagmatic settings can be reclaimed. If one accepts their

general conclusion, it is tempting to inquire as to how far this

hypothesis can be extended into other petrological realms. If

chemical variations of metabasalts retain information relating to

their genesis (tectonic setting), for example, this would be most

helpful in reconstructing the history of basalts from greenstone

belts.

Pearce and Cann (1) type diagrams are prepared by selecting

a training set for which the tectonic settings of all of the

analyses are known and obtaining a projection in which overlap of

the fields of the known groups is minimized. IF, the training set

is representative of a larger target population of interest, the

projection may allow assignment of an "unknown11 (an analysis not

part of the training set) to one of the recognized groups. As

the ratios of the variables are retained when percentages are

formed, the search for such fields presumes that there are limits

on the ratios of the three variables which identify a particular

tectonomagmatic setting. The selection of three components and

projection onto the plane of the ternary, however, does ignore

potentially useful information and one could argue that a

dimension-reducing procedure such as principal components

analysis might lead to a more satisfactory and potentially useful

display form.

However, a successful analysis of date with any multivariate

procedure requires more than an understanding of the procedure

itself. Additionally, the form of the data should be such that

statistical procedures can be rationally interpreted. The

subcomposition Ti-Zr-Y-Sr, for example, is part of a set of

percentages and therefore subject to all of the concerns

previously expressed by Chayes (2), Butler (3) and others

concerning difficulties in interpreting both statistical measures

of relationship (such as the correlation coefficient) and

empathetic analysis of "patterns and trends" expressed in some

compositional sub-space.

Simply stated, a set of composition percentages contains a

mix of information from at least two sourcesi

(1) physical/chemical relations among the variables

(2) a change in the structure of the data as a result of

a transformation such as percentage formation.

Statistical procedures typically allow one to recognize a

behavior pattern that departs from a hypothesized expected

behavior. The difficulty in interpreting percentages arises as a

result of the mix of information noted above. For example, given
,

12

TRACE ELEMENT DISTRIBUTIONS...

Butler, J.C.

.a statistically significant correlation between Zr and Ti, can

one automatically assume that contribution from the mechanical

process of forming percentages is negligible? Is it possible, in

fact, that the mechanical process is the only one operative for a

given measured relationship? Can the investigator separate these

two effects in a given situation and assess their influence?

Until recently (Aitchison, 4,5) these questions received a

great deal of discussion and warning (Chayes, 2 and Butler, 3)

but no defined solution. Aitchison (4,5) presents a set of

procedures that ultimately are designed to allow an invstigator

to make use of information contained within a set of percentages

and these procedures are adequately described in the literature.

A training set of average Ti-Zr-Y-Sr analyses of 35 modern

basalts (including 24 from Pearce and Cann (1) and 11 drawn from

the current literature) with known tectonic settings was drawn

from the literature Space is insufficient to tabulate these raw

data and details will be published elsewhere; copies of the raw

data, however, are available from the author).

Aitchison's tests for basis independence (4) and complete

subcompositional independence (5) both reject their respective

null hypotheses (6). Thus, the investigator is assured that the

mechanical contribution is not dominant and that a physicalchemical interpretation is warranted. Each analysis was

normalized to its geometric mean and eigenvalues and associated

eigenvectors extracted from the variance-covariance matrix of the

resulting log-row-centered data using principal components

analytical procedures. The first two eigenvalues account for

some 92% of the total variance and a plot of the first two

principal component scores is given in Figure 1. The boundaries

are empirical and constructed so as to isolate the known

tectonomagmatic groups. The distribution of scores successfully

delineates (1) the Nithin Plate Basalts, (2) the Ocean Floor

Basalts, and (3) the Arc Basalts. The principal component scores

are computed as followsi

Score 1 - -0.371*Ti-0.067*Zr-0.399*Y+0.836*Sr

Score 2 = -0.338*Ti-0.560*Zr+0.740*Y+0.158*Sr

where the individual variables are expressed in log-row-centered

form. In keeping with Pearce and Cann's suggestions (1), Ti is

defined as Ti02 times 100 and Y is defined as 3Y. As one is

dealing with a logarithmic function, multiplication by a constant

changes the scale of the resulting projection but not the spatial

relationships. Ten sets of analyses from the literature were cast

into the space defined in Figure 1. In general, the

tectonomagmatic settings predicted from Figure 1 are in excellent

agreement with interpretations by the respective authors. Of

prime concern in this case, however, is the effect of

metamorphism on such subcompositions. Many authors (1) have

noted that Sr is easily mobilized during low to intermediate

grade metamorphism whereas Ti, Zr and Y remain relatively

constant. Three of the 35 analyses are plotted in Figure 2 with

additions and subtractions of 10% and 30% total Sr. These sets

of points define sets of straight lines which are subparallel to

the X-axis. Note that the "trend" of these lines is such that it
Nithin Plate and Plate

the above model for Sr

r; i-f V ^ *> ^ *'

 f<

 . 13

ORIGINAL PAGE IS

OF POOR

TRACE ELEMENT DISTRIBUTIONS...

Butler, J.C.

may be possible to differentiate between

Margin Basalts (Ocean Floor plus Arc) if

mobilization holds.

Perhaps a combination of detailed knowledge of the geology

of a particular greenstone assemblage plus judicious use of

diagrams analogous to Figure 1 will enable the investigator to

see through effects which heretofore may have masked

petrogenetically significant information.

References! (1) Pearce, J.A. & Cann, J.R. (1973), Earth planet.

Sci. Lett..19.290. (2) Chayes, F., (1971) Ratio Correlation.

Univ. Chicago Press, Chicago, 111. (3) Butler, J.C. (1979) Amer.

Mineral.. 64. 1115. (4) Aitchison, J. (1981) Math. Geol.. 13.

175. (5) Aitchison, J. (1984) Math. Qeol.. 16. 617. (6) Moronow,

A. & Butler, J.C., (1986) Comp. and Qeos.. in press.

FIGURE 1. A plot of

the first two

principal components

for the training

set of 35 basalt

analyses. Numbers

refer to specific

analyses which are

available as a

separate from the

author. Boundary

curves are empirical

and drawn to

isolate the tectonomagmatic facies.

.65

.15

CM

UJ

oc

o

-.35

-.85

\ 15

\ •• ARC

OCEAN-FLOOR \ 17 '• Low-K 16

6

 34 ^\ ^ " ^ *--18-

• * .

7

^v 20 23 *|

911 •_&•- Calc-alkall 21

* • 13 *** . 33 " ~~~

• -* 24* 27

• • 34

26* 2*5 29.

WITHIN-PLATE

35 28

30

-.2 .8

SCORE 1

1.3 1.8

FIGURE 2. Additions

and subtractions of

10% and 30% Sr for

three of the basalt

analyses in the

space defined by the

first two principal

components. Note

that the trends are

parallel to each

other and subparallel to the

boundary between the

Mi thin Plate and

Plate Margin facies.

.65

CM

O

O

CO

.15

-.35

-.85

N,

OCEAN-FLOOR

\

ARC

-so — • 10

WITHIN-PLATE

-.2 .8

SCORE 1

1-3 1.8
14 N86-23095

BARBERTON GREENSTONE BELT VOLCANISM: ' SUCCESSION, STYLE, AND

PETROGENESIS; Gary R. Byerly and Donald R. Lowe, Department of

Geology, Louisiana State University, Baton Rouge, Louisiana 70803 USA

The Barberton Mountain Land is a small but remarkably wellpreserved and accessable early Archean greenstone belt along the

eastern margin of the Kaapvaal Craton of southern Africa. Although

there is some question about the role of structural repetition of

various units, detailed mapping in the southern portion of the belt

leads to the conclusion that a substantial thickness is due to original

deposition of volcanics and sediments (1). In the area mapped, a

minimum thickness of 12km of predominantly mafic and ultramafic

volcanics comprise the Komati, Hooggenoeg, and Kromberg Formations of

the Onverwacht Group, and at least one km of predominantly pyroclastic

and epiclastic sediments derived from dacitic volcanics comprise the

Fig Tree Group. Much greater apparent thicknesses of Fig Tree are due

to numerous structural repetitions. The essentially non-volcanic

Moodies Group lies conformably on top of the Fig Tree. The position

or correlation for the Sandspruit and Theespruit Formations relative to

the above units is not known. The Barberton greenstone belt formed

primarily by ultramafic to mafic volcanism on a shallow marine platform

which underwent little or no concurrent extension. Vents for this

igneous activity were probably of the non-constructional fissure type.

Dacitic volcanism occurred thoughout the sequence in minor amounts.

Large, constructional vent complexes were formed, and explosive

eruptions widely dispersed pyroclastic debris. Only in the final

stages of evolution of the belt did significant thrust-faulting occur,

generally after, though perhaps overlapping with, the final stage of

dacitic igneous activity.

The volcanic succession in the Barberton greenstone belt is often

used as a general model for greenstone belt stratigraphy (2). Previous

workers have suggested that volcanism there was cyclic, ultramafic to

mafic to felsic, on a scale that ranged from tens of meters to tens of

kilometers in stratigraphic thickness, with small cycles superimposed

on large cycles. In the grossest sense, the base of the sequence is

predominantly komatiitic and the top dacitic, but beyond this the

detail of volcanic succession is complex. Thin units of dacitic tuff

are recognized within the Komati Formation and komatiitic lavas are

interbedded with Fig Tree Group sediments. Simple, small-scale cycles

are not present. Sequences previously interpreted as small-scale

cycles are now known to represent thick, stratiform alteration zones of

mafic and ultramafic lavas to silicic rocks with a remarkably calcalkaline-like chemistry (3). Systematic increases in Si,K, and Rb

accompany decreases in Fe,Mg, and Ni, while Al,Ti, and Zr remain

constant from base to top in these cyclic units(4). Throughout these

alteration zones the flows typically have mafic volcanic textures and

structures, and are usually fine-grained and in places pillowed.

Preserved spinels in silicified rocks initially crystallized in mafic

or ultramafic lavas. After taking into account the nature of this

common style of alteration it appears there are no obvious systematic

trends in lava composition in the stratigraphic sequence. Notably,

however, the two thick sequences of dacitic volcanics seem to represent

prolonged volcanic episodes with no mafic or ultramafic activity.
BARBERTON VOLCANISM page two . 15

Byerly and Lowe

Komatiitic or basaltic igneous activity seems also to occur with

little or no other type of igneous activity in three or four thick

sequences.

Styles of igneous activity vary primarily as a function of lava

composition. Komatiites throughout the sequence occur as massive flows

with typical spinifex textures or as thick flows that often display

cumulus-layered bases or as pillowed flows and only rarely as

hyaloclastite units. In most sections the flows are quite thin,

typically 50cm to 5m, and only rarely up to 50m. They are rarely

vesicular, suggesting deep water extrusion, but in several sections

interbedded sediments are of shallow-water origin. We have observed no

vent complexes for the komatiites, though they are relatively

widespread units. The komatiitic unit within the Hooggenoeg can be

traced for over 50 km of strike around the Onverwacht anticline. The

komatiitic unit beneath the Msauli Chert crops out over a similar

distance. Only in the uppermost komatiitic unit is there a local

lateral facies; here the lavas interfinger with shallow marine

sediments and were of more local, extent, though again no vents are

recognized.. Basaltic igneous activity is characterized by thick to

thin lava flows, in places pillowed or massive and only rarely by

pillow breccias. Two, separate basaltic sequences in the Hooggenoeg

crop out for 50 km along the Onverwacht anticline. These lavas are

non-scoriaceous, but commonly contain up to 5% vesicles primarily as

radial vesicles about the margins of pillows. The lower basalts of the

Kromberg occur as a thick sequence of lapilli tuffs, especially thick

on the west limb of the Onverwacht anticline (5). These units are

locally crosscut by irregular dikes and sills of basalt, and in places

contain blocks and bombs of both juvenile and accidental lithologies.

They appear to represent near-vent facies and were perhaps similar to

modern basaltic cindercone fields. Some lithologies in this unit are

moderately scoriaceous. Laterally, these units are represented by

interbedded sediments and pillowed to massive lava flows. Dacitic

igneous activity is represented on two different scales: by the

relatively common tuffaceous units that occur throughout the section,

and by very thick sequences of lavas, pyroclastics, and epiclastics at

two locations in the sequence (1). Thin, typically a few tens of cm

but rarely to a few tens of m, tuffaceous units occur throughout the

sequence, and are usually completely altered to a micromosaic of quartz

and sericite. Textures are remarkably well preserved, however, and

indicate highly pumiceous particles often in the form of accretionary

lapilli commonly in graded airfall beds. These units are regionally

extensive, greater than SO km strike length, 'but associated vent

complexes are not found. The two major dacitic units, at the top of

the Hooggenoeg and top of the Fig Tree, clearly represent vent

complexes. They form complex associations of lava flows or domes,

breccias, and tuffs hundreds of m thick. Along strike systematic

changes in lithologies can be recognized where sedimentary rocks

represent debris being shed off the constructional vent complex.

These units do not appear to be laterally interbedded with more mafic

lavas.

Petrogenesis of Barberton greenstone belt volcanics is not likely

a single, one-stage process. Indeed, the succession of units and

common isolation of one compositional group from the others may even
16

BARBERTON VOLCANISM page three

Byerly and Lowe

require a separate petrogenesis for komatiites, basalts, and dacites.

Komatiites from the top and the base of the sequence are remarkably

similar in composition (4,6,7). They are typical of komatiites

worldwide except for very low Al/Ti, very high Ti/V, and other ratios

that require a very depleted upper mantle source (6). Otherwise, most

compositional variation within the komatiitic suite seems consistent

with low-pressure fractionation of olivine, later joined by

clinopyroxene in komatiitic basalts. Basalts of the Hooggenoeg and

Kromberg Formations have typical tholeiitic compositions, including a

pronounced iron-enrichment and lack of alumina-enrichment, that can be

produced by low-pressure f ractionation of plagioclase, clinopyroxene,

and olivine. Immobile trace elements and their ratios, such as very

low LREE/HREE, also require a depleted upper mantle source.

Compositional data are not inconsistent with a single liquid line of

descent of komatiites and basalts. While both komatiite and basalt

sequences suggest substantial low-pressure fractionation there is not

generally an adequate mass of layered intrusives to account for this

fractionation in situ. The Barberton sequence contains less than 5%

layered intrusives, yet basaltic komatiites and Fe-rich basalts each

require 50% or more fractional removal of crystalline phases from their

parental melts which must have taken place beneath the present level of

the greenstone belt. Dacites are the only intermediate to silicic

magmatic rocks found. They range from 60-70% Si02, 15-16% A1203, and

have Na2O/K20 ratios of about 3 in the freshest samples and are thus

trondjhemitic in character. They display extreme fractionation of LREE

to Y, and have very high concentrations of highly incompatible elements

such as U and Th. Plagioclase and hornblende are the major phenocryst

phases in all dacites. Some also contain either quartz or biotite as

phenocrysts. Their compositions suggest a source that was mafic in

composition and a relatively small degree of partial melting of an

assemblage dominated by amphibole. They are not related to associated

basalts by any simple, one-stage magmatic process, though could be

related to a second stage of igneous activity at the base of a thick,

hydrated pile of mafic volcanics.

REFERENCES

(1) Lowe et al,1985,Precamb.Res.,27,165-186.

(2) Anhauesser,1971,G.S.Aust.S.P.#3,57-70.

(3) Byerly et al,1983,LPSC XIV,84-85.

(4) Byerly and Lowe,1985,LPSC,XVI,101-102.

(5) Smith and Erlank,1982,in Komatiites,347-398.

(6) Viljoen et al,1983,in Contr.Geol.Barberton, 1-20.

(7) Ransom, 1986,M.S. Thesis,LSU.
N86-23096 17

TECTONIC SETTING AND EVOLUTION OF LATE ARCHEAN GREENSTONE

BELTS OF SUPERIOR PROVINCE, CANADA: K.D. Card, Geological Survey

of Canada

Late Archean (3.0-2.5 Ga) greenstone belts are a major component of the Superior

Province of the Canadian Shield where alternating, metavolcanic - rich and

metasedimentary - rich subprovinces form a prominent central striped region bordered in

part by high-grade gneiss subprovinces, the Pikiwitonei and Minto in the north, and the

Minnesota River Valley in the south. The high-grade gneiss subprovinces are

characterized by granulite facies gneiss of plutonic and supracrustal origin, and by

abundant plutonic rocks. Minnesota River Valley has rocks older than 3.5 Ga; absolute

ages of Pikiwitonei and Minto rocks re unknown but Minto does have north-south

structural trends distinctive from the dominant east-west structures of Superior

Province.

Volcano-plutonic subprovinces of'

Superior Province consist of generally narrow;

sinous greenstone belts bordered and intruded

by voluminous plutonic rocks, including

tonalitic gneiss, synvolcanic plutons, and

younger foliated to massive, generally

composite plutons, ranging from quartz

dorite to granite and syenite. Supracrustal

rocks of the greenstone belts include

komatiitic, tholeiitic, calc-alkalic, and rare

alkalic volcanics with volcanogenic clastic

(wacke, conglomerate) and chemical (iron

formation, chert) sediments. Most,

greenstone belts consist of several lensoid/

overlapping piles each on the order of 100 km

in maximum dimension and approximately 5

to 10 km thick and commonly comprising!

several volcanic cycles. Some cycles consist of a lower komatiitic -tholeiitic basalt

sequence, a middle tholeiitic basalt -andesite sequence, and an upper calc-alkalic daciterhyolite-andesite sequence. Other cyles are bimodal tholeiitic basalt-dacite (rhyolite)

sequences. Minor alkalic and shoshonitic volcanics and associated alluvial/fluvial

sediments are present in some belts where they unconformably overlie older volcanics

and synvolcanic plutons. In term of rock types, sequences, and overall configuration,

many Superior Province greenstone belts are closely comparable to modern island arcs.

Superior Province greenstone belts typically have upright folds with curved,

bifurcating axial surfaces, steep foliations and lineations, and major domal culminations

and depressions, the products of polyphase deformation. Some belts, however, display

low angle foliations and faults, overturned sequences, and recumbent and downward -

facing structures suggestive of thrust-nappe style tectonics (1,9,10).

The enclosing gneissic and plutonic rocks display domal structural patterns, again

the product of polyphase deformation involving recumbent folding and diapirism.

Metamorphic grade in the greenstone belts is generally subgreenschist to greenschist in

the central parts grading outward to low pressure amphibolite facies in belt margins and

surrounding plutonic gneisses.

*.•»' Proleroioic Plwn.roxoic locks

—•-• Subfxownc. boundwy

UtCHEAN SUBPflOWNCE TYPE
'/ Vs

18 TECTONIC SETTING AND EVOLUTION OF LATE ARCHEAN

K.D. Card

The contacts between greenstone belts and enclosing plutonic rocks, and between

the greenstone-rich subprovinces and adjacent plutonic subprovinces are generally either

intrusive or tectonic. An unconformity between greenstones and older granitoid rocks

has been demonstrated only at Steeprock, Ontario(5) and although younger volcanics and

older plutonic rocks are juxtaposed in a number of places, faults, mylonites, or shear

zones invariably intervene. Dextral transcurrent faults trending EW and NW and sinistral

faults trending NE form subprovince boundaries in part, as do NE and EW trending

thrusts. One notable product of this faulting, the Kapuskasing Structural Zone, exposes

granulites considered to represent upthrusted lower crust (7,8). Late alkalic volcanic -

fluviatile sediment sequences are spatially related to major transcurrent faults and may

represent deposition in pull-apart basins formed by alternating periods of transtension

and transpression in strike-slip zones.

Interpretation of geophysical data

shows changes in depth to the Conrad

Discontinuity and to the Moho from one

subprovince to another, indicating

significant structural relief across their

faulted boundaries (*). Greenstone belts of

Abitibi and Wabigoon subprovinces generally

extend to depths of only 5 to 10 km(3)

whereas metasedimentary gneisses of

English River Subprovince and plutonic

rocks of Winnipeg River Subprovince may

extend to depths of 10 to 20 km(4).

Juxtaposition of high-pressure granulites of

the Kapuskasing zone with low-pressure

greenschist-amphibolite facies rocks of

Abitibi Subprovince implies structural relief

of 15 to 20 km across the boundary thrust

(7,8).

Metasedimentary

(English River, Quetico,

subprovinces

Pontiac etc.)

LEGEND

Fault

Tnrual twit

Tranacurraflt lautt

Normal fault

Province boundary

consist mainly of turbidite wacke and pelite metamorphosed at grades ranging from low

greenschist at belt margins to upper amphibolite and locally, low-pressure granulite in

belt interiors. Anatectic, s-type granitic rocks are prevalent in the migmatitic, highgrade interiors of the metasedimentary belts.

Most metasedimentary subprovinces have a linear aspect attributable to

transcurrent boundary faults and isoclinal folds with subhorizontal to subertical axes,

late structures superimposed on earlier complex, recumbent folds and dome-basin

structures. In areas where contacts between metasedimentary and volcano-plutonic

subprovinces are unfaulted, there appear to be rapid facies transitions from sedimentary

to dominantly volcanic sequences. Preliminary isotopic age data also indicate that the

sedimentary and volcanic sequences of some adjacent subprovinces are broadly coeval.

U-Pb zircon dates demonstrate that volcanic, plutonic, deformational, and

metamorphic events of relatively brief duration affected large parts of Superior

Province and that there are detectable differences in ages of these events from one area

to another(6). In the northwest (Sachigo, Berens, Uchi subprovinces) major volcanism and
ORIGINAL PAGE IS

OF POOR QUALITY

.4"

TECfofjIC^SETTING AND EVOLUTION OF LATE ARCHEAN

CARD, K.D.

accompanying plutonism occurred at about

3.0 to 2.9 Ga, 2.85 to 2.80 Ga, and 2.75 to

2.7 Ga. These voclanic episodes were

followed by major deformation,

metamorphism, and plutonism about 2.73 to

2.7 Ga. In the south (Wabigoon, Wawa,

Abitibi subprovinces) volcanism and

plutonism occurred mainly between 2.75 and

2.69 Ga, followed by major deformation,

metamorphism, and plutonism at about 2.70

to 2.66 Ga. There is evidence for somewhat

younger (2.65 to 2.63 Ga) metamorphicplutonic events, or of later closure of

isotopic systems, in the high-grade rocks of

the metasedimentary belts and of the

Kapuskasing zone.

In summary, Superior Province

consists mainly of Late Archean

supracrustal and plutonic rocks with Middle

Archean gneisses in the south and possibly in

the north. The Late Archean supracrustal

sequences are possibly mainly of island-arc and inter-arc affinity, although continental

rift zone settings have also been postulated(2). Abundant plutonic rocks include early

synvolcamc intrusions and later synorogenic and post-orogenic intrusions derived in part

from the mantle and in part from crustal melting caused by thermal blanketing of newlythickened continental crust combined with high mantle heat flux.

The contemporaneity of magmatic and deformational events along the lengths of

the belts, coupled with the structural evidence of major compression and transcurrent

faulting, is consistent with a subduction-dominated tectonic regime for assembly of the

Superior Province orogen. Successive lateral and vertical accretion of volcanic arcs and

related sedimentary accumulations, accompanied and followed by voluminous plutonism,

resulted in multi-stage crustal thickening and stabilization of the Superior craton prior

to emplacement of mafic dyke swarms and Early Proterozoic marginal rifting.

REFERENCES

1. Attoh, K. (1981); Geological Survey of Canada, Paper 81-1B, p. *9-5«f.

2. Ayres, L.D. and Thurston, P.C. (1985): Geological Association of Canada, Special

Pa n«*r /fin^/i/i^Q/ l * *^ *-•»**

reaper £O) p* <2*TlT"Oov/«

3. Gupta, J.K., Thurston, P.C. and Dusanowskyj, T.H. (1982): Precambrian Research,

v. 16, p. 233-255.

*' Hfji

 D*H

*

 and Brisbin> w

-

c

- <1982)' Candian Journal of Earth Sciences, v. 19, p.20<*9-

19

5. Joliffe, A.W. (1955): Economic Geology, v. 50, p. 373-398.

6. Krogh, T., Corfu, F. and Davis, D.W. (1985): G.A.C. Spec. Paper 28, p. 358.

7. Percival, J.A. and Card, K.D. (1983): Geology, v. II, p. 323-326.

8. Percival, J.A. and Card, K.D. (1985): G.A.C. Spec. Paper 28, p. 179-192.

9. Poulsen, K.H., Borradaile, G.J., and Kehlenbeck, M.M. (1980): Canadian Journal of

Earth Sciences, v. 17, p. 1358-1369.

lO.Thurston, P.C. and Breaks, F.W. (1978): Geological Survey of Canada, Paper 78-10,

p. "9—62.

ORIGINAL PAGE OS

OF POOR QUALITY
20 N86-23097

HIGH PRECISION U-PB GEOCHRONOLOGY AND IMPLICATIONS FOR THE TECTONIC

EVOLUTION OF THE SUPERIOR PROVINCE; D.W. Davis, F. Corfu, and T.E. Krogh,

Oack Satterly Geochronology Laboratory, Royal Ontario Museum, 100 Queen's

Park, Toronto, Ontario. MBS 2C6

The underlying mechanisms of Archean tectonics and the degree to which

modern plate tectonic models are applicable early in earth's history

continue to be a subject of considerable debate. A precise knowledge of the

timing of geological events is of the utmost importance in studying this

problem. The high precision U-Pb method has been applied in recent years to

rock units in many areas of the Superior Province. Most of these data have

precisions of about ± 2-3 Ma. The resulting detailed chronologies of local

igneous development and the regional age relationships furnish tight

constraints on any Archean tectonic model.

Superior province terrains can be classified into 3 types:

1) low grade areas dominated by meta-volcanic rocks (greenstone belts).

2) high grade, largely metaplutonic areas with abundant orthogneiss and

foliated to massive I-type granitoid bodies.

3) high grade areas with abundant metasediments, paragneiss and S-type

plutons.

Most of the U-Pb age determinations have been done on type 1 terrains

with very few having been done in type 3 terrains.

A compilation of over 120 ages indicates that the major part of igneous

activity took place in the period 2760-2670 Ma, known as the Kenoran event.

This event was ubiquitous throughout the Superior Province.

There is, however, abundant evidence for the widespread occurrence of

pre-Kenoran volcanoes and sialic crust, especially north of the WabigoonEnglish River subprovince boundary. In the Uchi and Sachigo subprovinces

there are volcanic periods about 3000-2900 Ma (2,3) and 2850-2800 Ma (2,3)

in age which underlie the Kenoran sequence. The Kenoran rocks are in part

disconformable on the older sequences. The general absence of angular

unconformities along with other evidence such as the presence of mature

sandstones in the Sachigo subprovince (4), implies an extended period of

crustal stability preceding the Kenoran event. Tonalites 2950-3200 Ma in

age are found in the Favourable Lake (5), North Spirit Lake (3) and Winnipeg

River Belts (6,7,8), suggesting a pre-Kenoran crust-forming event at about

3000 Ma. Evidence for the existence of an extensive pre-Kenoran continent

is especially strong in the Winnipeg River Belt, a type 2 terrain. Recent

data obtained from a type 2 terrain in the Wabigoon subprovince also

indicates 3000 Ma volcanic and plutonic sequences (9). This indicates that

type 2 terrains in many cases include pre-Kenoran crust, and that pre-Kenoran

crustal material may be locally present in type 2 areas throughout the

Superior Province.

The earliest Kenoran magmatism consisted of eruption of tholeiitic

basalt platforms. These are difficult to date but in some areas pre-date

2750 Ma (10). Intermediate-felsic calc-alkaline volcanistn occupied a time

span from about 2750-2700 Ma and led to the construction of large composite

volcanoes. The transition to calc-alkaline volcanism was associated with the

emplacement of layered basic intrusions and contemporaneous tonalitegranodiorite plutons, without major deformation. This was followed by the

intrusion of high alumina trondhjemite-granodiorite plutons, in some cases

accompanied by later calc-alkaline volcanism. This resulted in the development of large intravolcanic batholiths (11). Significant regional
U-PB GEOCHRON. & TEC. EVOL. OF SUPERIOR PROV.

Davis, D.W. et al. 21

deformation began relatively late, at about 2700 Ma over much of the Superior

Province. In some areas, such as the Wabigoon greenstone belt, it

significantly post-dated the bulk of calc-alkaline intrusive and volcanic

activity (12).

Some greenstone belts underwent at least two periods of deformation.

A Dl event affected the calc-alkaline sequences and pre-dated sedimentation

and eruption of alkaline volcanic rocks (e.g. Tamiskaming sequence). The

Tamiskaming-type sequences were then affected by a later deformational (D2)

event. The ages of the late sequences and the D2 event are bracketed

between 2695 and 2685 Ma in the southern Wabigoon and Shebandowan subprovinces (13), but may have been 15-20 Ma earlier in the Oxford Lake belt in

the northern Superior Province.

The causes of regional deformation are unclear. It may have been partly

the result of diapiric remobilization of the intravolcanic batholiths (e.g.

Wabigoon greenstone belt), accompanied by regional compression, perhaps due

to the intrusion of marginal late granitoid plutons (e.g. Batchawana belt)

(14). The presence of nappe structures in some areas such as the Winnipeg

River belt (15) and the southern Wabigoon subprovince (16) further

complicates the tectonic picture. The final expression of strain was the

establishment of large strike-slip faults, which often separate type 1 from

type 2 and type 3 terrains.

The deformational event was accompanied by intrusion of late tectonic

plutons, most of which have ages in the range 2700-2670 Ma (17). This

resulted in cratonization and brought the Kenoran event to an end. Locally,

single volcanic centers passed through this cycle from initial volcanism to

terminal deformation in time spans as short as 30 Ma (18).

Although there is some indication of a secular younging in the peak of

Kenoran igneous activity in a N-S direction, the broad simultaneity and short

time spans of crustal events argue against any simple model for growth of the

Superior Province by accretion of island arcs (19). Furthermore, there is a

strong vertical control on magmatic and metamorphic ages. The oldest Kenoran

plutons occur high in the crust while the youngest plutons and metamorphic

ages are found at deeper crustal levels in more uplifted and eroded terrains

such as the Berens River subprovince, parts of the Winnipeg River belt and

the Kapuskasing structural zone (20).

Despite considerable work on the felsic units in type 1 greenstone

terrains, there is almost no evidence of inherited zircon components derived

from significantly older sialic material. The intravolcanic granitoid rocks

and thick felsic volcanic sequences were largely derived by differentiation

processes from mafic precursors within the period of Kenoran activity (11).

However, greenstone belts evidently,did develop adjacent to older sialic

blocks. Evidence for this, found in the Wabigoon greenstone belt, includes

pre-Kenoran granitoid clasts in a conglomerate marginal to the belt and the

existence of marginal unconformities between 3000 Ma tonalite in the Winnipeg

River belt and Kenoran volcanic and plutonic sequences (7). Abundant mafic

dykes intrude the older units below these unconformities and indicate a

tensional stress regime.

The bulk of the evidence presently available argues for a model, in which

greenstone belts were initiated by rifting of older sialic crust and the

formation of narrow ocean basins. The fault controlled nature of many

subprovince boundaries as well as the fact that volcanism was at times nearly

coeval throughout the Superior Province suggests that rifting may have been

concentrated along major early lithospheric breaks.
U-PB GEOCHRON. & TEC. EVOL. OF SUPERIOR PROV.

22 Davis, D:W. et al.

Evidence for subduction in late Archean tectonic processes is missing.

The absence of an effective subduction mechanism would have inhibited ocean

spreading. If the intracratonic rifts were not able to open into wide ocean

basins they would have been reworked in place, undergoing dominantly vertical

tectonic processes. Continued mantle-derived mafic magmatism may have led to

thickening and differentiation of the crust to produce the large amounts of

calc-alkaline material now present in type 1 terrains.

Any model for tectonic development can only be tentative and subject to

the constraints of a constantly expanding data set. Some of the major

questions remaining for geochronology are the extent in time and space of

pre-Kenoran material and its deformational history and the origin and

basement of the metasedimentary belts. These questions can only be resolved

by much more extensive work in type 2 and type 3 terrains. .

References

1 Kroner A. (ed.) (1981) Precambrian Plate Tectonics, Elsevier, Amsterdam.

2 Nunes P.O. and Thurston P.C. (1980) Can. J. Earth Sci. J7, p. 710-721.

3 Corfu F., Wallace H., Wood J. and Ayres L.D. (1984) (abstract) Geol.

Assoc. Canada annual meeting, London, Ont., p. 54.

4 Donaldson J.A. and Ojakangas R.W. (1977) Can. J. Earth Sci. 14.

p. 1980-1990.

5 Corfu F. and Ayres L.D. (1984) Contrib. Mineral. Petrol. 88, p. 86-101.

6 Krogh T.E., Harris N.B. and Davis G.L. (1976) Can. J. Earth" Sci. 2»

p. 1212-1215.

7 Clark G.S., Bald R. and Ayres L.D. (1981) Can. J. Earth Sci. J.8,

p. 94-102.

8 Corfu F, Beakhouse G.P., Stott G.M. and Sutcliffe R.H. (1985) (abstract)

Institute on Lake Superior Geology, 31st annual meeting, Kenora, Ont.,

p. 17.

9 Davis D.W. and Jackson M.C. (1985) Ontario Geological Survey Summary

of Field Work.

10 Davis D.W., Blackburn C.E. and Krogh T.E. (1982) Can. J. Earth Sci. ^9,

p. 254-266.

11 Davis D.W. and Edwards G.R. (1985) Summary of Research, Ontario

Geoscience Research Grant Program 1984-1985.

12 Davis D.W. and Edwards G.R. (1982) Can. J. Earth Sci. 19, p. 1235-1245.

13 Corfu F. and Stott G.M. (1985) (abstract) Institute on~Take Superior

Geology, 31st annual meeting, Kenora, Ont., p.. 15.

14 Corfu F. and Grunsky E.C. (1984) (abstract) Geol. Assoc. Canada annual

meeting, London, Ont., p. 54.

15 Beakhouse G.P., Stott G.M. and Sutcliffe R.H. (1983) Ontario Geological

Survey Summary of Field Work, MP116, p. 5-10.

16 Poulsen K.H., Borradaile G.J. and Kehlenbeck, M.M. (1980) Can. J. Earth

Sci. 17, p. 1358-1369.

17 Krogh T7E., Davis D.W., Nunes P.O. and Corfu F. (1982) (abstract) Geol.

Assoc. Canada annual meeting, Winnipeg, Man., p. 61.

18 Davis D.W. and Edwards G.R. (1985) (abstract) Institute on Lake Superior

Geology, 31st annual meeting, Kenora, Ont., p. 20.

19 Langford F.F. and Morin J.A. (1976) Amer. J. Sci. 276, p. 1023-1034.

20 Percival J.A. and Krogh T.E. (1983) Can. J. Earth 5cT. 20, p. 830-843.
N86-23098

CRUSTAL STRUCTURE OF THE ARCHEAN GRANITEGREENSTONE TERRANE IN THE NORTHERN PORTION

OF THE KAAPVAAL CRATON

J.H. de Beer (National Physical Research Laboratory, CSIR, P.O.

Box 395, Pretoria 0001, S. Africa), E.H. Stettler (Geological

Survey, P.O. Box X112, Pretoria 0001, S. Africa), J.M. Barton Jr,

D.D. van Reenen and J.R. Vearncombe (Dept. of Geology, Rand

Afrikaans University, P.O. Box 524, Johannesburg 2000, S.Africa).

Recent investigations of the electrical resistivity, gravity

and aeromagnetic signatures of the various granite-greenstone

units in the northern portion of the Kaapvaal craton have revealed three features of significance: 1) the Archean greenstone belts are shallow features, rarely exceeding 5 km in depth;

2) the high resistivity upper crustal layer typical of the lower

grade granite-greenstone terranes is absent in the granulite facies terrane and 3) the aeromagnetic lineation patterns allow

the granite-greenstone terrane to be subdivided into geologically

recognisable tectono-metamorphic domains on the basis of lineation frequency and direction.

In the Pietersburg, Sutherland and Murchison greenstone

belts geoelectrical investigations showed that the greenstone

lithologies have a lower resistivity than the surrounding granitic terranes. Positive gravity anomalies over the greenstone

belts are related to more dense metamorphosed ultramafic and

mafic rocks in the belts compared to surrounding granitic rocks.

Numerical modelling of the geophysical data indicates that the

greenstone belts are asymmetrical structures, being thicker

along the southeastern flanks. The belts are underlain by high

resistivity, low density granitic rocks of which two types are

distinguished by their average densities: a lower density series

(density = 2600 kg m~^) corresponding to 2650 Ma granodioritic

plutons and a higher density series (density = 2670 kg m~3) comprising the older gneissic terrane. The younger series is well

developed along the southern margins of the greenstone belts and

occurs locally along the northern margins. Primary layering and

tectonic fabric within the greenstone lithologies are subvertical. Thicknesses measured across layering exceed the depth of
CRUSTAL STRUCTURE NORTHERN KAAPVAAL CRATON

24 De Beer, J.H. et al.

the belts, suggesting no simple rotation of the greenstone lithologies but instead a truncation at shallow depths of structurally repeated (folded and imbricated) greenstone belts. This

truncation may be a .major recumbent deformation zone, recumbent

syntectonic granite or a late intrusive contact.

Deep resistivity soundings indicate significant changes in

the regional structure of the crust in the northern portion of

the Kaapvaal craton corresponding to changes in metamorphic

grade and tectonic style. ,In the low-grade granite-greenstone

terrane the upper 10 km or less of the crust is characterized by

high-resistivity rocks (approximately 100 000 ohm.m) overlying a

more conductive layer (approximately 5 000 ohm.m) to a depth of

about 35 km.. Below this possible mantle rocks with a resistivity

of about 50 ohm.m.occur. Where the granulite facies rocks of

the northernmost Kaapvaal craton (southern marginal zone of the

Limpopo belt) occur the approximately 100 000 ohm.m layer is

absent and rocks with a resistivity of 5 000 ohm.m extend to a

maximum depth of 35 km, where they overlie possible mantle rocks.

The significance of these variations in the physical properties

of .the Kaapvaal craton will be addressed.

The aeromagnetic lineation pattern in the study area can be

divided into distinct domains on the basis of the lineation

frequency and direction. Although these magnetic anomalies are

due to mafic and ultramafic dykes they reflect an inherent fabric in the crust. The domain boundaries correspond to known

tectonic and/or metamorphic transitions. One such boundary

being the orthoamphibole rehydration isograd that marks the

transition between the granulite.facies terrane of the southern

marginal zone of the Limpopo belt (northern Kaapvaal' craton) and

the lower-grade rocks to the south. It is clear that the lineation pattern does not reflect the different lithological units

in the area and the Sutherland and Pietersburg greenstone belts

are, for example, not reflected in the aeromagnetic lineation

pattern. This suggests that they are an internal component of

certain domains and not marking domain sutures.
N86-23099

EXTENSIONAL TECTONICS DURING THE IGNEOUS EMPLACEMENT OF THE

MAFIC-ULTRAMAFIC ROCKS OF THE BARBERTON GREENSTONE BELT. M.J. de Wit, Lunar and

Planetary Institute, 3303 NASA Road One, Houston, TX 77058 and BPI Geophysics,

University of the Witwatersrand, Johannesburg, South Africa.

The simatic rocks (Onverwacht Group) of the Barberton greenstone belt,

which occur in at least 3 regional thrust nappes, are part of the Jamestown

ophiolite complex*. This ophiolite, together with it's thick sedimentary

cover (Fig Tree and Moodies Groups) occupies a complex thrust belt. Field

studies have identified two types of early faults which are entirely confined to

the simatic rocks and are deformed by the later thrusts and associated folds.

The first type of fault (Fla) is regional and always occurs in the simatic rocks

along and parallel to the lower contacts of the ophiolite-related cherts (Middle

Marker and equivalent layers; for their distribution see Fig. 1, de Wit et al.,

this volume). These faults zones have previously been referred to both as

flaser-banded gneisses2 and as weathering horizons3. (Fla) zones consist of

anastomosing, cross-cutting and folded extension veins which have internal

cross-fibrous growth textures. Vein filling minerals are predominantly calcite,

less often quartz. The veins are separated by schistose to proto-mylonitic

folia of fuchsite, chlorite, sericite and serpentines (Fig. 1). In general the

zones range between l-30m in thickness. The veins formed by a succession of

dilation-diffusion increments4'

5 and subsequently deformed during simple shear

to form banded gneisses (Fig. 1; in this poster presentation, polished slabs of

these rocks will be displayed). The simatic host rocks close to (Fla) zones,

are ubiquitously brecciated and extensively altered (carbonatized and/or

silicified) as documented by the major elements, stable and radiogenic isotope

compositions (REE are relatively stable). This alteration is related to a

extensive hydrothermal-fluid/rock interaction. It has been postulated that the

dilatancy-anisotropy of the fault zones was related to a hydraulic

fracturing-gliding mechanism in a geothermal environment". Episodic decrease

of fluid overpressure due to movement in these zones would cause boiling,

calcite precipitation and crack-sealing with a concomittant resistance to

movement of the cherty cap-rock". Displacements along these zones are

difficult to estimate, but may be in the order of 1-102 km. The structures

indicate that the faults formed close to horizontal, during extensional shear

and were therefore low angle normal faults. In many areas, both the faults and

their overlying cherts, are cut by subvertical simatic intrusions of the

Onverwacht Group (Fig. 2). Thus (Fla) zones overlap in age with the formation

of the ophiolite complex. The second type of faults (Fib) are vertical

brittle-ductile shear zones, which crosscut the complex at variable angles and

cannot always be traced from plutonic to overlying extrusive (pillowed) simatic

rocks. (Fib) zones are therefore also apparently of penecontemporaneous origin

with the intrusive-extrusive igneous processes (Fig. 2). Thus (Fib) zones may

either represent transform fault-type activity or represent root zones

(steepened extensions) of (Fla) zones. Both fault types indicate extensive

deformation in the rocks of the greenstone belt prior to compressional

overthrust tectonics, and at least (Fla) implies regional extensional tectonics

and probably block rotation during the formation of the ophiolites.

References

(1) de Wit et al. this volume. (2) de Wit (1982) Journal of Structural Geology,

4, p. 117-136.^3) Lowe et al. (1985) Precambrian Research, 27. p. 165-186.

T4) Ramsay, J.G. (1980) The crack-seal mechanism of rock deformation. Nature,

284, p. 135-139. (5) Ramsay, J.G. and Huber, M.I. (1983) Strain analysis.

Academic Press. (6) de Wit et al_., Economic Geol.. 77^ p. 1783-1802.
EXTENSIONS TECTONICS IN THE BARBERTON BELT

I . •-

de Wit, M. J.

ORIGINAL PAGE IS

OF POOR QUALITY

Fib

FIG.2

Figure Captions: Fig. 1 (a) Anastomosing/crosscutting carbonate extension veins

(pale-grey) with thin schistose folia (dark grey). Sections up to 30 meters

thick entirely composed of this rock-type constitute flaser-banded

tectonites . (b) as In (a), showing the cross-fibrous carbonate growth

textures in the veins. Different shades of.grey are due to variations in

concentration of inclusion bands and trails . (c) Internal brecciation and

shearing of cross-fibres (vertical) yielding (subhorizontal) protoraylonites.

(d) gneissose-mylonitic fabric following shearing and flattening of extension

veins. (2) Block diagram of area near the Onverwacht bend (see Fig. 1. de Wit

this volume for location - as outlined by the box marked Fig. 2) showing the

disposition of both (Fla and Fib) fault zones. Note how vertical metawehrlite

intrusions (grey) have cross-cut and incorporated screens of middle marker-like

cherts underlain by (Fla) gneissose tectonites (shear zones).
N86-23100

27

A MID-ARCHEAN OPHIOLITE COMPLEX, BARBERTON MOUNTAIN LAND; M.J. de Wit,

Lunar and Planetary Institute, 3303 NASA Rd. One, Houston, TX 77058 and BPI

Geophysics, University of the Witwatersrand. Roger Hart, School of Oceangraphy

OSU, Corvallis, Oregon 97331. Rodger Hart, SC Nuclear Sciences, University of

the Witwatersrand, Johannesburg, South Africa.

New field observations and structurally restored geologic sections through

the southern part of 3.5-3.6 Ga Barberton greenstone belt (Fig. 1) show that

it's mafic to ultramafic rocks form a pseudostratigraphy comparable to that of

Phanerozoic ophiolites; we refer to this ancient ophiolite as the Jamestown

ophiolite complex*. It consists of an (in part sheeted, Fig. 2)

intrusive-extrusive mafic-ultramafic section, underlain by a high-temperature

tectono-metamorphic residual peridotitic base, and is capped by a chert-shale

sequence which it locally intrudes. Geochemical data support an ophiolitic

comparison (Fig. 3). Fractionation of high temperature melting PGE's (>

2500°C) in the residual rocks suggest a lower mantle origin for the precursors

this crust^. An oceanic rather than arc-related crustal section can be

inferred from the absence of contemporaneous andesites. This ancient simatic

crust was thin (<3 km), contains a large ultramafic component («25%), is

pervasively hydrated (> 95%) with H2o contents ranging between 1-15% and

consequently has a low density (r=2.67 g/cmj)-3.

The entire simatic section has also been chemically altered during its

formation by hydrothermal interaction with the Archaean hydrosphere (Fig. 4).

Only an igneous "ghost" major element geochemistry is preserved. This

regionally open-system metasomatism may have increased the MgO content of the

igneous rocks by as much as 15%. The most primitive parent liquids, from which

the extrusive sequence evolved, may have been "picritic" in character. Rocks

with a komatiitic chemistry may have been derived during crystal accumulation

from picrite-crystal mushes (predominantly olivine-clinopyroxene) and/or by

metasomatism during one or more subsequent episodes of hydration-dehydration

(Fig. 5).

The Jamestown ophiolite complex provides the oldest record with evidence

for the formation of oceanic lithosphere at constructive tectonic boundaries.

Our observations are in agreement with models predicting higher oceanic Archean

heat flux per unit ridge length than today, associated with deep mantle diapiric

upflow. Because of its low density, this ophiolite resisted subduction during

subsequent tectonism; it was obducted to form part of a thrust complex^.

References

(1) de Wit, M.J., Hart, R. and Hart, R. (1985) The Jamestown ophiolite

complex, Barberton Mountain Belt: a reconaissance study of a composite section

through 3.5 Ga simatic lithosphere (under review). (2) Tredoux, M., de Wit,

M.J., Hart, R., Armstrong, R., Sellschop, J.P.E., (1985) PGE in the 3.6 Bon

Accord Ni-Fe-Co-Cr occurrence, Barberton South Africa: implications for

crust-mantle-core segregation. 4th |ntern. Platinum Symposium. Toronto, August,

1985, abstract; also de Wit et_ aT7~(1985), EOS, Nov. 12, p. 1114. (3) de Wit

M.J. (1985) What the oldest rocks say, (abstract) Workshop on The Earth as a

Planet, LPI and GSA, Orlando, Florida, Oct. 27.
28

.

A MID ARCHEAN OPHIOLITE COMPLEX

de Wit, M. J. et. al. PAGE IS

POOR QUALHY

PIG!

METMttOFHIC ZONES

-|«TM^R1 VOLCANICS

PILLOW LAVAS

SHALLOW LEVEL

INTHUSIVES

DIKES AXO SILLS

a

F/G5

AMPHIBOLITES

ULTBA«AF5CS

TECTW1T6S

v-Hfl

1^ — Ustsr

''
A MID ARCHEAN OPHIOLITE COMPLEX

de Wit, M. J. et. al.

29

IF16.5 Figure Captions: (1) Simplified geological map of an area In the southern part

of the Barberton greenstone belt studied between 1978 and 1985. (2) Vertical

sheeted Intruslves with pale chilled margins from a 30 meter river outcrop

(exposed during 1984 draught) at locality A, Fig. 1. Note the remnant chert

xenollth (a; arrow), (b) clearly depicts the splitting 1n two of an earlier

Intrusion (1) by a later one (2). (3) Representation of 1LOQ (a) of the

Barberton rocks (black) plotted In their restored pseudo-strati graphic sequence

compared to Phanerozolc ophlolltes and oceanic crust (open symbols) (b) REE data

from Barberton; this plot compares favourably with Phanerozolc ophlolltes and

oceanic crust. (4) Binary correlation plots of HgO, CaO, 510, and H.,0 for

rocks of oceanic crust (open symbols) and from the Jamestown ophlollte complex

(closed symbols). These plots Illustrate the close correlation between the

major oxides concentrations and the degree of hydratlon 1n these environments.

For comparison, the slopes of the chemical flux 1n the Galapages hot spring

fluids are also shown. (5) This figure shows that the bulk rock MgO/MgO + Feo

ratio of Barberton Komatlltes are enriched 1n HgO over that of the original

melt. The enrichment may be the result of either crystal accumulation or

magnesium metasomatism during hydrothermal alteration; there Is textural

evidence that both mechanism were Important. At any rate the plot clearly shows

that the HgO composition of the silicate liquids which formed the Barberton

Komatllte was between that of Gorgona Island (15-222 HgO) and Alexo (28% HgO),

and may have been of plcrltlc composition. All diagrams from reference 1.
30

FELSIC IGNEOUS ROCKS WITHIN THE BARBERTON GREENSTONE BELT: HIGH

CRUSTAL LEVEL EQUIVALENTS OF THE SURROUNDING TONALITE-TRONDHJEMITE TERRAIN,

EMPLACEO DURING THRUSTING. M.O. de Wit, Lunar and Planetary Institute, 3303

NASA Road One, Houston, TX 77058 and BPI Geophysics, University of the

Witwatersrand, Johannesburg. A.H. Wilson, University of Natal,

Pietesmaritzburg, South Africa.

Felsic rocks within the 3530 + 50 myrs1 simatic rocks of the Onverwacht

Group of the Barberton greenstone Felt have traditionally been mapped as

recurring volcanic units within a continuous stratigraphic succession. In the

past, these felsic units have been interpreted to be part of several mafic to

felsic volcanic cycles within this sequence. Some of these silicic layers have

been shown to be silicified simatic rocks2*

3. Our field data (Fig. 1)

indicates that the genuine felsic igneous rocks are predominantly shallow level

intrusives and subsurface felsic domes associated with only minor volcanics and

volcanoclastics. A 3.360 +_ 1 myrs (U-Pb, zircon)4 age from the main felsic

intrusion indicates that its emplacement post-dated the simatic rocks of this

greenstone belt between 120-220 myrs. Our geochemical results also show that

the felsic igneous rocks are not directly related to the mafic-ultramafic rocks

of the Onverwacht Group. On the contrary the major trace and REE data (Fig. 2)

all indicate that these felsic units are high-level equivalents of the

widespread, and time-equivalent, trondhjemite-tonalite plutons which either

intrude the lower parts of the greenstone belt, or with which they are in

tectonic contact.

Structural and stratigraphic analysis indicates that the felsic intrusions

were emplaced along thrusts during sedimentation and a prolonged period of

horizontal compressional stress exerted on the greenstone belt (Fig. 3). Thus,

integrated, the data suggest that the simatic rocks of the Barberton greenstone

belt were thrust across an actively stoping plutonic environment and that the

greenstone belt is at least partly allochthonous (Fig. 4).

References

(1) Hamilton, P.J. et al. (1979) Sm-Nd dating of the Onverwacht Group volcanics,

South Africa. Nature. 279. p. 298-300. (2) de Wit ejt al. (1982) Archean

abiogenic and probable biogenic structures associated with mineralized

hydrothermal vent systems and regional metasomatism with implications for

greenstone belt studies. Econ. Geol.. 77, p. 1783-1802. (3) Hanor, J.S. and

Duchac, K.C. (1985) Geochemical consequences of the metasomatic conversion of an

early Archean Komatiite sequence into chert. GSA Abstracts, Geol. Annual

Meeting, Orlando, p. 603. (4) Van Niekerk, C.B. and Burger, A.J. (1969) A note

on the mininum age of the acid lava of the onverwacht series of the Swaziland

system. Trans. Geol. Soc. South Africa. 72. p. 9-21.
FELSIC IGNEOUS ROCKS EMPLACED DURING THRUSTING

de Wit, M. J..and Wilson, A. H. 31

ORIGINAL PAGE IS

OF POOR QUALITY

F1G.1

•:—«!**' _!^v. em,*,

—4—iiw

'•i"! ..'»-'.»-'« «H

'i'".. <M«-r«

34 MOST ALTLUD

1C* HO Sn.6jOon.Dy Ef n Ui

C| M Figure Captions: (1) Simplified geological map of the southern part of the

rl(>..d Barberton greenstone belt, showing location of main silicic (felslc) rocks. (2)

(a) Statistical analysis of major element data from the felslc Igneous rocks

within the study area, compared to those of the surrounding tonallte and

trondhjemltes. The felslc Igneous rocks are clearly divided Into two groups (I

and II) 1n which both extrusive (ex) and Intrusive (In) samples are represented.

The two groups are geochemlcally similar to the trondhjemltes (thin frequency

boxes) and tonal1tes (bold frequency boxes) (b) Chondrlte normalized REE

patterns of Intrusive and extrusive representatives of both groups of felslc

Igneous rocks from within the greenstone belt, compared to the granitoid plutons

surrounding the greenstone belt.
32

FELSIC IGNEOUS.ROCKS EMPLACED DURING THRUSTING

de Wit, M. J. and Wilson, A. H.

FIG-3

TECTOMC TRANSPORT

O

UPUFT-EWJSON

O

ORIGINAL PAGE IS

OF POOR QUALfTY

TECTOMC TRANSPORT DIRECTION (Dj)

PLUTONS

I STENTOR

HNELSPRUTT

• KAAPVAUEY

ETNELSHOOGTE

H STOLZBERG

SI THEESPRUIT

5DDALMEIN

FELStC ROCKS TO FELSIC VENT

>*! TECTONIC

^ TRANSPORT

TONAUTE-TROMfiHIEUrTC UEU8 FROM TECTOMCMLY BUHED GREENSTONE AUjOCHTHONS

(3) Schematic representation of the

tectonic-intrusive emplacement of the felsic igneous rocks as composite sills

close to the Interface between the mafic-ultramafic (simatic) rocks of the

Onverwacht rocks (diagonal lines) and the overlying Fig Tree-like silts-shales.

Note how the lower contacts of the sills are predominantly tectonic (thrusts)

whilst the upper contacts are predominantly preserved igneous contacts. (4)

Plan and section of the Barberton greenstone belt (black) and the surrounding

granitoid terrain (white). The map shows the generalized 0, tectonic

transport directions, the felsic Igneous rocks internal to the greenstone belt,

and the gnelssose fabric in the surrounding tonalite-trondhjemite plutons. Note

that large scale sloping of the greenstone belt by the surrounding and intruding

granitoids is suggested by the outcrop pattern of the felsic igneous rocks

(eg. compare this pattern to the shape and outline of the Stentor pluton). The

section schematically shows the lower parts (3-5 km) of the greenstone belt

thrust over the granitoid terrain whilst the latter syntectonically Intrudes and

engulves the greenstone belt: this process 1s thought to have formed

recumbent-like mantle-gneiss folds (probably sheath-like 1n 3-diment1ons).

Regional disruption and sloping of the greenstone belt occurs during Intrusion

of Na-rich felsic phases from the plutons of the granitoid terrain Into the

greenstone belts, along thrusts generated during the tectonic emplacement of the

entire greenstone belt. The section represents a restoration prior to

subsequent horizontal flattening which later deformed and rotated the rock units

and their contacts into a pseudo-synformal structure. All diagrams from de Wit,

Wilson and Armstrong (1985 under review).
N86-23101

33

HEAT FLOW AND HEAT GENERATION IN GREENSTONE BELTS; Malcolm J.

Drury, Earth Physics Branch, Energy, Mines and Resources Canada, 1

Observatory Crescent, Ottawa, Ontario, K1A OY3 Canada

Heat flow has been measured in Precambrian shields in both greenstone

belts and crystalline terrains. Values are generally low, reflecting the

great age and tectonic stability of the shields; they range typically

between 30 and 50 mH/m2, although extreme values of 18 and 79

mW/m2 have been reported (1,2). For large areas of the earth's

surface that are presumed to have been subjected to a common

thermotectonic event, plots of heat flow against heat generation appear

to be linear (3,4), although there may be considerable scatter in the

data. The relationship is expressed as:

Q = Q0 + D A0 [1]

in which Q is the observed heat flow, A,, is the measured heat

generation at the surface, Q0 is the "reduced" heat flow from the

lower crust and mantle, and D, which has the dimension of length,

represents a scale depth for the distribution of radiogenic elements.

Most authors have not used data from greenstone belts in attempting to .

define the relationship within shields, considering them unrepresentative

and preferring to use data from relatively homogeneous crystalline rocks,

e.g. (5).

The heat generated by radioactive decay is expected to be less in

basic than in acidic rocks because of their different chemistry. Hence

we would expect heat flow in greenstone belts to be lower than that in

adjoining crystalline areas if the greenstones are thick, but to be

similar if the belts are merely superficial. Table 1 is a compilation of

data from seven Precambrian shields. Only those data specifically

identified as being from greenstone belts, or those for which geological

descriptions are unambiguous, are used in column 2. There is the

possiblity that some of the data identified as being from crystalline

areas are in fact from greenstone belts.

Table 1. Compilation of heat flow data for Precambrian shields, listed

according to-geological setting. The ratio in column 4 is that of the

mean heat flow in the greenstone belts to that in crystalline areas of

the shield.

Shield Mean and 1 s.d. heat flow (mW/m2)

All sites

Canadian*3

Canadian*1*

Baltic0

W. African*

Indian*

Australian^

42±8

43±10

40±6

36±12

64±15

40±8

(22)

(10)

(26)

(19)

( 6)

(16)

Greenstones

39±5

40±9

41±6

35

44

38±8

(8)

(6)

(4)

(1)

(1)

(8)

Crystalline

43±10

48111

40±6

38±11

68±12

42±8

(14)

( 4)

(22)

(18)

( 5)

( 8)

Ratio

0.91

0.83

1.03

0.92

0.65

0.90

Brazilians 52±11 (12) 51±18 (2) 52±10 (10) 0.98

a Superior province, reference 6 with additional data not yet published;

b Churchill province, 1; c 5,7,8,9,10,11; d 12; e 2; f 13, 14,

15; S 16. * - heat flow values adjusted for glaciation effects.
HEAT FLOW AND HEAT GENERATION IN GREENS. BELT

34

DRURY. M.

Although it appears from column 4 of Table 1 that mean heat flow in

greenstone belts is indeed lower than that in crystalline areas of the

shields, there is, in all shields except one, considerable overlap of the

two values. The exception is the Indian Shield, but there is only one

value from a greenstone belt for that. Further, in most cases no

statistical signficance can be inferred as there are fewer data for

greenstone belts than for crystalline areas. Taking the mean values, the

heat flow from crystalline areas is apparently approximately 10% higher

than that from greenstone belts.

Not all heat flow data used for compiling Table 1 had associated heat

generation data. The most complete set is for the Canadian shield

(Superior and Churchill provinces). Linear least squares regression for

those data yields:

Q0 D r Q* A* n

(mW/m2) km (mH/m2) (pW/m3)

Greenstone belts 33±4 7±6 0.45 37±7 0.51±0.46 7

Crystalline areas 26±6 12±4 0.67 40±9 1.16*0.51 11

where n is the number of data pairs, r is the correlation coefficient, Q*

is the mean heat flow and A* is the mean heat generation of borehole

samples. The correlations are low and statistically the differences

between the parameters for the two crustal types are insignificant.

However, assuming that radiogenic elements are distributed uniformly with

depth to D, the value of D for the greenstones suggests that they are

approximately 7 km thick, a value compatible with those cited by Condie

(17). The data also suggest that the heat flow - heat generation

relationship for the greenstones could be written as

Q - Qo = CDc-Dg)Ac + DgAg [2]

in which subscripts g and c refer to greenstone and crystalline crust and

Q0 is the reduced heat flow for the crystalline crust. This can be

seen by inserting appropriate.values for greenstones and crystalline

terrain into equation [2]. It implies that the greenstones are underlain

by normal crystalline crust, including 5 km of upper crust, but that they

are not allochthonous, replacing 7 km of that crust rather than simply

overlying it.
,"' •;• ' * ^ HEAT FLOW AND HEAT GENERATION IN GREENS. BELTS

35

DRURY, M,

2 Rao R.U.M., Rao G.V. and Naraiii H. (1976) Earth Plan. Sci. Lett.. 30,

p. 57-64.

3 Roy R.F., Blackwell D.D. and Birch F. (1968) Earth Plan. Sei. Lett..

5, p. 1-12.

4 Lachenbruch A.E. (1968) J. Geophys. Res. 73_, p. 6977-6989.

5 Swanberg C.A., Chessman M.D., Simmons G., Smithson S.B., Gronlie G. and

Heier K.S. (1974) Tectonophysics. 23, p. 31-48.

6 Jessop A.M. and Lewis T.J. (1978) Tectonophysics. 50. p. 55-77

7 Puranen M., Jarviamki P., Hamaleinen U. and Lehtinen S. (1968)

Geoexploration. 6_, p. 151-162

8 Lubimova E.A., Karus E.V., Firsov F.V., Starikova G.N., Vlasov V.K.,

Lyusova L.N. and Koperbakh E.B. (1972) Geothermics. 1, p. 81-89

9 Parasnis D.S. (1975) Geophys. J. Roy. Astr. Soc.. 43. p. 531-554

10 Parasnis D.S. (1982) Geophys. J. Roy. Astr. Soc.. 71, p. 545-554

11 Gordienko V.V., Zavgorodnya O.V., Moiseenko U.I. and Smyslov A.A.

(1985) Geophys. Res. Lett.. 12., p. 549-552

12 Brigaud F., Lucazeau F., Ly, S. and Sauvage, J.F. (1985) Geophys. Res.

Lett.. 12, p. 549-552

13 Howard L.E. and Sass J.H. (1964) J. Geophys. Res.. 69, p. 1617-1625

14 Hyndman R.D., Lambert I.B., Heier K.S., Jaeger J.C. and Ringwood A.E.

(1968) Phys. Earth Plan. Int.. 1, p. 129-135

15 Jaeger J.C. (1970) Earth Plan. Sci. Lett.. 8, p. 285-292

16 Viterello I., Hamza V.M. and Pollack H.N. (1980) J. Geophys. Res.. 85_,

p. 3778-3788.

17 Condie, K. (1981) Developments in Precambrian geology 3, Elsevier

Scientific Publishing Co.. New York.
36

KOMATIITE GENESIS IN THE ARCHAEN MANTLE, WITH IMPLICATIONS FOR

THE TECTONICS OF ARCHAEN GREENSTONE BELTS; D. Elthon, Lunar and Planetary

Institute, 3303 NASA Rd.l, Houston, TX 77058 and Department of Geosciences,

University of Houston — University Park, Houston, TX 77004.

The presence of ultramafic lavas (komatiites) associated with Archaen

greenstone belts has been suggested to indicate very high increments (50-

80%) of partial melting of the Archean mantle [e.g., 1-3], Such extensive

melting of the Earth's mantle during the Archaen might have profound

effects on the early tectonic and chemical evolution of the planet [e.g., 4

& 5], although problems associated with keeping the komatiite liquid in

equilibrium with the residual mantle at such high increments of melting has

cast doubt upon aspects of extensive melting [e.g.,6 & 7]. Two important

aspects of the origin of koraatiites are discussed below.

JL WHAT IS THE NATURE OF PRIMARY KOMATIITE LIQUIDS?

One of the most fundamental aspects of understanding the tectonic and

geochemical mode of origin for komatiites is the problem of komatiite

primary magmas. The identification of primary komatiite magmas is

complicated by the extensive metamorphism that these rocks have typically

undergone and by olivine (+ minor spinel) crystallization at low pressures

( 1 atm). The crystallization of olivine rapidly depletes a komatiite

liquid in MgO, such that the most likely candidates for primary magmas are

those with the highest MgO contents.

Previous efforts to evaluate primary komatiitic liquids have proposed

that they might contain as much as 33% MgO [2] or 30% MgO [8]. These

studies have relied principally on comparison of the compositions of

olivines crystallized in high-pressure experimental studies of komatiites

with relict olivines found in komatiites (as high as Fo94).

The Fe-Mg exchange between olivine and basaltic-komatiitic liquids has

recently been summarized by [9], in which they present equations for

calculation of olivine-liquid equilibria over a wide temperature (1074-

1600° C.) and pressure (1 atm to 25 kbar) range. The KD values for a wide

range of komatiites (>20% MgO) were calculated using this equation and

range from 0.28 to 0.31 at temperatures of 1450-1650°C at 1 atm. This

olivine-liquid equilibrium is shown in Fig. 1 along with the compositions

of the most magnesian olivines in komatiites (olivine and koraatiite

compositions from [8] and references therein). The 1 atm KD values have

been used here because the present author considers it most likely that the

olivines in komatiites have crystallized at very low pressures ( 1 atm);

previous investigators [2 & 4] have used KD values from high pressure

experiments, which are substantially higher [10 & 11].

The data shown in Fig. 1 (horizontal lines connect the olivine

compositions with the liquid from which they could have crystallized)

indicate that the komatiite olivines probably have crystallized from

liquids with Fe/Mg >0.230. This Fe/Mg (0.230) corresponds to 22-25% MgO in

the komatiites, depending upon the FeO content of the liquid. This data

indicates that the most magnesian olivines in komatiites could have

crystallized from liquids with 22-25% MgO, in contrast to previous

estimates of 30-33% MgO. These liquids will have liquidus temperatures of

1500°C at 1 atm pressure. More MgO-rich komatiites have probably become

enriched in MgO as a consequence of olivine accumulation and/or Mg

metasomatism.
KOMATIITES

D. Elthon 37

.Fig. 1

Cation Fe/Mg

in olivine

versus cation

Fe/Mg in the

liquid. The

horizontal

lines connect

the composition

of olivines

in komatiites

with liquids

in equilibrium

with these

olivines.

.12

£ .08

O)

2

04

CD

O

7839

49J

V.O

Fo

- 90

-62

-94

.10 .20

Cation Fe/Mg in Uiq.

II. WHAT PERCENTAGE OF MELTING IS REQUIRED TO PRODUCE KOMATIITES?

As noted above, it is generally assumed that a very high increment

of melting (50-80%) is required in order to generate komatiites from the

Earth's mantle. Experimental studies of the melting of reasonable mantle

compositions have shown that very magnesium-rich magmas may be produced at

high increments of melting [e.g., 11-13]. As outlined below, however,

these. MgO-rich magmas produced by very large increments (40-80%) of melting

within the mantle are NOT komatiites.

A pseudo-liquidus phase diagram for evaluating the petrogenesis of

komatiites is shown in Fig. 2. At low pressures ( 1 atm), some magmas

(those above the 01-R join) will crystallize augite as the first pyroxene

and others (those below the OL-R join) will crystallize pigeonite or

orthopyroxene first. Field and petrographic studies of komatiites have

shown that they crystallize augite as the first pyroxene in virtually all

instances. Most terrestrial magmas also crystallize augite as the first

pyroxene; boninites are an obvious exception. Also shown in Fig. 2 is the

field for the compositions of komatiites from Munro Township, which

crystallize augite as the first pyroxene [14].

A partially schematic melting path for melting of the mantle is shown

in Fig. 2 for melting at 15 kbars. At small to moderate increments of

melting «~30%), the primary liquids will lie above the OL-R join, but will

lie below the OL-R join at larger increments of melting. The extent of

melting required to produce primary magmas below the OL-R join will vary as

a function of the composition of the mantle and the pressure of melting,

but it is clear that high increments of melting that might produce dunite

or OPX-poor harzburgite residues will produce primary magmas that will lie

below the OL-R join and will evolve to crystallize orthopyroxene and/or

pigeonite before augite. The extent of melting most likely to produce

komatiitic magmas is more like 20-25% rather than the 50-80% previously

proposed. Although not discussed by the previous authors, this feature is

further apparent in the data of [12 & 13],

Spinels from komatiites have Cr/(Cr+Al) from 0.70 to 0.80 [15],

which would suggest a slightly higher increment of partial melting of the

mantle that occurs in the present-day suboceanic mantle [16], rather than

the much higher increments proposed in previous studies.
38

KOMATIITES

D. Elthon

Fig. 2

Pseudo-liquidus phase diagram

projected from plagioclase[17].

Pseudo-invariant points at

1 atm (R), 10, 15, 20

and 25 kbars are

shown.

CPX

In summary, it is suggested that the extent of partial melting that

produces koraatiite primary magmas is 20-25% and that these magmas have 22-

25% MgO or less. This substantially lower estimate for the extent of

melting and eruption temperatures will certainly influence those tectonic

characteristics of greenstone belts associated with the dynamics of mantle

upwelling and convection.

REFERENCES:

[I] Viljoen, R.P. and Viljoen, M.J. (1969) Geol. Soc. S. Afr.. Spec.

Publ. 2, 275-296.

[2] Green, D.H., Nichols, I.A., Viljoen, M.J. and Viljoen, R.P. (1975)

Geology 3, 11-15.

[3] Bickle, M.J., Hawkesworth, C.J., Martin, A., Nisbet, E.G. and O'Nions,

R.K. (1976) Nature 263, 577-580.

[4] Bickle, M.J. (1978) Earth Planet. Sci. Lett. 40. 301-315.

[5] Nisbet, E.G. and Walker, D. (1982) Earth Planet. Sci. Lett. 60.

105-113.

[6] Arndt, N.T. (1977) Contrib. Mineral. Petrol. 64. 205-221.

[7] Waff, H.S. (1980) J. Geophys. Res. 85. 1815-1825.

[8] Bickle, M.J. (1982) in: Komatiites N.T.Arndt & E.G.Nisbet (eds).

pp.479-494.

[9] Ford, C.E., Russell, D.G., Craven, J.A. and Fisk, M.R. (1983) J._

Petrol. 24, 256-265.

[10] Grover, J.E., Lindsley, D.H. and Bence, A.E. (1980) Proc, Lunar Sci.

Conf. 11. 179-196.

[II] Mysen, B.O. and Kushiro, I. (1977) Am. Mineral. 62. 843-865.

[12] Jaques, A.L. and Green, D.H.(1980) Contrib. Mineral. Petrol. 73. 287-

310.

[13] Takahashi, E. and Kushiro, I. (1983) Am. Mineral. 68, 859-879.

[14] Arndt, N.T., Naldrett, A.J. and Pyke, D.R. (1977) J. Petrol. 18. 319-

369.

[15] Cameron, W.E. and Nisbet, E.G. (1982) in: Komatiites N.T.Arndt & E.G.

Nisbet (eds.) pp. 29-50.

[16] Dick, H.J.B. and Bullen, T. (1984) Contrib. Mineral. Petrol. 86. 54-

76.

[17] Elthon, D. and Scarfe, C.M. (1984) Am. Mineral. 69. 1-15.
N86-23103

39

THE YILGARN CRATON WESTERN AUSTRALIA : A TECTONIC

SYNTHESIS : I R.E.P. Fripp, Western Australian Institute of Technology

The Yilgarn Craton in Western Australia is one of the larger contiguous

preserved Archaean crustal fragments, with an area of about 650,000 square

kilometres. Of this, by area, about 70% is granitoid and 30% greenstone.

The Craton is defined by the Darling Fault on its western margin, by

Proterozoic deformation belts on its southern and northwestern margins, and

by unconformable younger sediments on its eastern and northeastern margins.

A regional geotectonic synthesis at a scale of 1:500,000 is being prepared.

This is based largely upon the 1:250,000 scale mapping of the Geological

Survey of Western Australia together with interpretation using geophysical

data, mainly airborne magnetic surveys.

On a regional basis the granitoids are classified as pre-, syn- and posttectonic (1) with respect to greenstone belt deformation. The post-tectonic

granitoids yield Rb-Sr isochrons of about 2.6 b.y., close to Rb-Sr ages for

the greenstones themselves which are up to about 2.8. b.y. old (2), although

data for the latter is sparse.

Contacts between earlier granitoids and greenstones which are not obscured

by the post-tectonic granitoids are most commonly tectonic contacts, intensely

deformed and with mylonitic fabrics. The general concensus however is that

there is a pre-tectonic, pre-greenstone sialic gneiss preserved in places

(1,3).

Existing models for the evolution of the belts involve 3 large basinal

structures ("broad elongate downwarps"), of which the Eastern one (the

Noreseman-Wiluna Belt) is considered to be a rift fault-bounded graben (1).

The postulated basins are separated by large tabular belts of discordant

post-tectonic granite when viewed regionally. This may be a 'red herring1.

It is possible that, for example, the entire greenstone package preserved on

the Craton was part of one basin, or numerous combinations and parts of

basins. There is no compelling diagnostic evidence collated to date to

postulate on the original disposition, geometry and relationships between

belts.

This synthesis is a preliminary attempt at addressing this problem, by

attempting to decipher the broad tectonic-strati graphic sequences preserved

and thereby to reconstruct, as far as is possible, the original nature of the

greenstones. There is structural evidence to suggest that the deformation

histories of the greenstones and some of their surrounding and occluded

granitoids involves early fold-nappe tectonics in places, and possibly thrust

nappes, as well as late large-scale imbrication or slicing. During early

deformation of the belts, massif-style nappe tectonics may have occurred in

places, on scales not dissimilar to those seen in young fold belts.

It is intended, with future work, to test these postulates and to examine

whether the tectonic history of the Yilgarn Craton is indicative of the loss

of considerable greenstone (back to the womb?) and perversely (sic), its

local preservation by obduction and stacking. How well can we reconstruct the

deformed granitoids and greenstones, in their undamaged state?
f ^ ; fi ''- .. < v ;/•<

40

 : THE YILGARN C^RATON WESTERN AUSTRALIA

Fripp, R.E.P.

REFERENCES

1. Gee, R.D., Baxter, J.L., Wilde, S.A. and Williams, I.R. (1981) Spec.

Publs geol. Soc. Aust., T_> p. 43-56.

2. de Laeter, J.R., Libby, W.G. and Trendall, A.F. (1981) Spec. Publs

geol. Soc. Aust., 7_, p. 147-157.

3. Archibald, N.J. and Bettenay, L.F. (1977) Earth planet. Sci. Lett.,

33_, p. 370-378 .
N86-23104

41

TECTONICS OP SOME AMAZONIAN GREENSTONE BELTS; Allan K. Gibbs,

Institute for the Study of the Continents, and Dept. Geol. Sci., Cornell

University, Ithaca, New York, 14853

Greenstone belts exposed amid gneisses, granitoid rocks, and less

abundant granulites along the northern and eastern margins of the Amazonian

Craton yield Trans-Amazonican metamorphic ages of 2.0-2.1 Ga. (Regional

geology: 1-13)« Early Proterozoic belts in the northern region probably

originated as ensimatic island arc complexes. The Archean Carajas belt in the

southeeastern craton probably formed in an extensional basin on older

continental basement. That basement contains older Archean belts with pillow

basalts and komatiites. Belts of ultramafic rocks warrant investigation as

possible ophiolites.

NORTHERN.BELTS - Volcanic rocks of the northern belts were erupted in the

Early Proterozoic (2.3-2.1 Ga)(l4-17). The contiguous belts of Guyana (18,19)

and Venezuela (20,21) closely resemble those of Suriname (7-9,22) and French

Guiana (1,4,16,23), though the two regions are separated by the Central Guiana

Granulite Belt. Typical sections consist of a lower flow and pillowed low-K

basalt-gabbro unit, overlain by interbedded mafic, intermediate, and felsic

volcanics of both tholeiitic and calc-alkaline suites; overlain by and

interstratified with volcaniclastic greywackes, pelites, and chemical

sedimentary rocks. Basalts with pronounced iron-enrichment and others with

high magnesium contents are both present, as are both tholeiitic and

calc-alkaline andesites and felsic volcanics (18,19,22,24). Generally

conformable tuffaceous and epiclastic conglomerates, greywackes, lithic

arenites, and shales appear petrographically and geochemically to have been

derived from the associated volcanic rocks, without significant contributions

from continental sources (18,25). The relative abundances and types of

volcanic and sedimentary rocks vary: felsic volcanics are irregularly

distributed, and magnesian basalts and possible komatiites are particularly

common in central French Guiana (22). Ultramafic, mafic, and anorthositic

intrusive complexes may be genetically associated with some of the volcanic

rocks (1,18,23). Some belts are overlain by quartz-rich epiclastic

sedimentary rocks that were folded and metamorphosed with the belts but appear

to be unconformable (1,13).

The northern belts have randomly-branching synclinal map patterns.

Prominent metamorphic foliations generally correlate with the regional folds,

with foliations locally crenulated or destroyed by younger shear deformation,

which elongated (WNW-ESE) both the belts and associated granitoid rocks.

Metamorphic grades range from amphibolite on the belts' peripheries to lower

greenschist and zeolite in the interiors. Diverse local mineral assemblages

indicate high, intermediate, and low-pressure metamorphic aeries. Anatectic,

two-mica granites intrude metapelitic schists along the northern periphery.

No evidence has been reported of basement-cover relations between the

northern belts and adjacent gneisses. Field observations and geochemical

similarities suggest that the greenstones pass into the intervening gneisses

by increase in metamorphic grade (14,15,17,26-28). The associated granulites

also appear to represent Early Proterozoic, rather than Archean crust

(16,27,29). Sm-Nd and Rb-Sr isotopic systematics indicate that little if any

older continental crust was involved in this greenstone-belt volcanism.

The northern belts are thought to have been originally contiguous with

the Birimmian belts of west Africa. Mature sedimentary rocks overlying the

greenstone belts have much in common with the Tarkwaian of West Africa.
42 AMAZONIAN GREENSTONE BELTS

Gibbs, A.K.

EASTERN BELTS - Belts of the east-central craton (30,31) have not been

adequately dated. Most lithostratigraphic sections have not yet been

resolved, in part due to intense deformation and common medium grade

metamorphism. Prominent banded iron formations, ultramafic schists, and

current-bedded, fuchsite-bearing quartz arenites and conglomerates are

present: these lithologies are uncommon in the northern belts. Small

enclaves of iron formations and chromite-bearing ultramafic rocks occur in

south and central Suriname, and might correlate with the east-central belts.

Archean greenstone belts with pillow basalts and komatiites, and belts of

serpentinite occur amid granitoid rocks and gneisses in the southeastern

craton, apparently forming a basement to the Serra dos Carajas belt (32). The

latter has a dominantly mafic bimodal volcanic suite, roughly 4-6 km thick and

dated at 2.75 Ga, overlain by 100-300 m of iron formation, and a 1-2 km thick

fine clastic and chemical sedimentary complex (33,34). The mafic rocks are

unlike typical Archean basalts and basaltic andesites, but have chemical and

isotopic evidence of contamination with older continental crust, like many

basalts of modern contiental extensional settings.

ACKNOWLEDGMENTS - Research in the Amazonian Craton has been substantially supported by the Guyana

Geological Survey, the U.S. National Science Foundation (under Grants No. EAR-7617026, 8207422,

and 8410379), and Utah International, Ltd. Generous institutional cooperation of DOCEGEO, the

United Nations Revolving Fund for Mineral Exploration, COGEKA; the BRGM, the Suriname Geology and

Mining Service, and ICOMI are also gratefully acknowledged.

REFERENCES h 1 Choubert, B. (1974) Mem. B.R.G.M., 81, 213 pp. [2] Martin Bellizzia, C. (1974)

Mem. Novena Conf. Geol. Inter-Guayanas, Venez. Min. Minas Hidroc. Bol. Geol. Publ. Esp.

6, pp. 251-305- [3l Mendoza, V., (1974) Segundo Congr. Latinoamer. Geol. Caracas, V. Ill, pp.

2237-2270. [4] B.R.G.M., (1979) Inventaire rainier du departement de la Guyane. [5] Choudhuri, A.

(1980) Precamb. Res., 13: 363-374. [6] Gibbs A.K., and Barren, C.N. (1983) Episodes 6, No. 2,

7-14- [7] De Vletter, D.R. (1984), Contrib. Geol. Suriname 8: 11-30. [s] Bosnia, W.,

Kroonenberg, S.B., van Lissa, R.V., Maas, K., and de Roever, E.W.F. (1983) Geol. Mijnb.

62:241-254. [9] Bosma, W., Kroonenberg, S.B., Van Lissa, R.V., Maas, K., and de Roever, E.W.F.

(1984) Cont. Geol. Suriname, 8: 31-82. [lO] Santos, J.O.S., (and Santos and S.O.C. Loguercio)

(1984) in Schobbenhaus et al. (eds.) Geologia do Brasil, DNPM, Brasilia, p. 59-91, 93-127. [ll]

Amaral, G. (1984) in F.F.M. de Almeida and Y. Hasui, 0 Precambriano do Brasil, Edgard Blucher,

Sao Paulo, p. 6-35. [12] Gibbs, A.K. , and Wirth, K.R. (1985) Workshop on Early Cruatal

Evolution, Lunar Planet. Sci. Inst. p. 87-91. [l3] Gibbs, A.K. (1985 submitted) 7th Geol.

Soc. Africa Conf. Botswana, November, 1985, 21 p. [l4l Gibbs, A.K., and Olszewski, W.J. , Jr.

(1982) Precamb. Res. 17: 199-214. [l5] Klar, G. (1979) Ph.D. Thesis, Case Western Reserve

Univ., 163 pp. [16] Gruau, G., Martin, H, Leveque, B., Capdevila, R. and Marot, A. (1985)

Precamb. Res. 30: 63-80. [171 Teixeira, W., Ojima, S.K., and Kawashita, K. (1984) D.N.P.M., II

Symp. Amazonico, Manaus, April 8-12, 1984, Anais, p. 75-86. [18] Gibbs, A.K. (1980) Ph.D.

Thesis, Harvard Univ., 385 pp. [19] Gibbs, A.K. , O'Day, P.A., and Renne^, R. (1984) Anais II

Amazonian Symp. Manaus, Brazil, April 8-15, 1984. D.N.P.M., Brazil p. 39-52. [20] Menendez, A.,

Benaim, N., and Espejo, A. (1974) Mem. IX Conf. Geol. Interguayanas, Ciudad Guayana, Venezuela,

1972, pp. 339-342.. [21] Benaim, N. (1974) IX Conf. Geol. Inter-Guayanas, Publ. Espec. Min.

Min. Hidroc. Venezuela, no. 6, pp. 198-206. [22] Marot, A., et al. (1984) 10 Reunion Ann.

Sci. Terre, Bordeaux, 1984, Soc. Geol. Fr. Edit. Paris. [23] Veenstra, E. (1983) Contrib.

lieoi. Suriname, 7, 134 p. [24J Renner, H. (1985) M.Sc. Diss. Cornell Univ., 252 p. [25]

O'Day, P.A. (1984) M.Sc. Diss., Cornell Univ., 154 p. [26] Cannon, R.T. (1964) Geol Surv.

British Guiana Bull 35, 83 p. [27] Priem, H.N.A., et al. (1978) in Zartman, R.E., (ed.), U.S.

Geol. Survey Open File Report 78-701, pp. 341-343. [28] Cordani, U.G., Teixeira, W., and Basei,

M.A.S. (1975) Anais 10th Inter-Guianas Geol. Conf., Belem, Brazil. Bol., no. 2, pp. 53-54

(aljair'.. ?9 3en Othman, 1)., Folve, M., and Allegre, C.J. (1984) Nature, 307: 510-515. [30]

Jorge Joao, X., and Marinho, P.A. (1982) Anais I Simp. Geol. Amazonia, Belem, v. 2, pp.

207-228. [31] Jorge Joao, X., et al. (1979) Geologia, 10, Secao Geol. Basica N. 7, DNPM

Brazil. 125 p. 1.32J Hirata, W.K. , Rigon, J.C., Kadekaru, K., Cordeiro, A.A.C., and Meireles, E.A.

(1982) Anais Simp. Geol. Amazonia I, 100-110. [33] Gibbs, A.K., Wirth, K.R., Hirata, W.K., and

Olszewski, Win. J. Jr. (1985 submitted) II Simposio Geol. Amazonia, Soc. Bras. Geol. Nucl.

Norte. Dec. 1-9, 1985, Belem. 17 p. [34] Wirth, K.R., Gibbs, A.K., and Olszewski, Wm. J. Jr.

(1985 submitted) II Simp. Geol. Amazonia, Soc. Bras. Geol. Nucl. Norte, Dec. 1-9, 1985

Belem. 12 p. [35] Cohen, H. (1985) Unpubl. Geology B.A. diss., Cornell Univ. 50 p. [36]

Onstott, T.C., Hargraves, R.B., York, D., and Hall, C. (1984) Geol. Soc. Amer. Bull

95:1045-1054. L37J Cordani, U.G., Tassinari, C.C.G., and Kawashita, K. (1984) Ciencias da Terra,

Institute for the Study of the Continents, Contribution No. 38
; £ -j - -

:

•^ % 'ii'** '«. ', * -•> *

AMAZONIAN GREENSTONE BELTS

Gibbs, A. K.

ORIGINAL PAGE IS

OF POOR QUALITY 43

Portions of the Amazonian and west African

cratona adjusted for postulated displacements

along Pan-African and older fault zones.

- Archean Imataca and Liberian

terranes

'r-Vr ~ Central Guiana Granulite Belt

*" ' <2.5 Ga

•• - greenstone belts (sensu lato)

- quartz-rich metasedimentary

rocks considered unconfofinable

on the greenstone belts

- granitoid rocks and gneisses

- areas with abundant continental

igneous and sedimentary cover

<1.9 Ga

Postions of the cratons (13, based in part on

35) is compatable with paleomagnetic data

(36), and juxtaposes geolgical features in

the two cratons. A major geological province

boundary must be present in the Amazonian

craton between the Carajas belt and the

northern belts (17,37,12): one possible

position is shown.
N86-23105

44

Spatial greenstone-gneiss relationships:

evidence from mafic—ultramafic xenolith

distribution patterns

A.Y. Glikson

Division of Petrology and Geochemistry, Australian

Bureau of Mineral Resources, Geology and Geophysics

ABSTRACT

The distribution patterns of mafic-ultramafic xenoliths

within Archaean orthogneiss terrain furnish an essential

key for the elucidation of granite-greenstone relations.

A complete gradation in scale exists between synclines,

large-scale outliers and outcrop-scale xenoliths of mafic

and ultramafic metavolcanic rocks. Accordingly, most

greenstone belts constitute "mega-xenoliths" rather than

primary basin structures. Transition along strike and

across strike between stratigraphically low greenstone

sequences and xenolith chains demonstrate their contemporaneity, as shown for example in Fig. 1 where the relationships between the Holenarsipur greenstone belts and

associated xenoliths in southern India are portrayed.

Regional to mesoscopic-scale characteristics of xenolith

swarms and their relations with early greenstone units

are well expressed in parts of the Pilbara Block, Western

Australia. Xenolith distribution patterns in domearcuate syncline gneiss-greenstone terrains define subsidiary gneiss domes within the batholiths. These terrains

represent least deformed cratonic "islands" within an

otherwise penetratively foliated deformed gneiss-greenstone crust. The oval gneiss domes are thought to have

developed originally by magmatic diapirism - evidenced by

intrusive relations and contact aureoles - followed by

late—stage solid state uprise related to isostatic adjustments. The late vertical movements were associated with

development of major shear zones along tectonized boundary zones of batholiths, where interdigitated deformed

gneiss—amphibolite schist intercalations were derived by

the attenuation of xenolith-rich orthogneiss. The deformation process involved interthrusting and refolding of

the interleaved plutonic and supracrustal units. The

exposure of high grade metamorphic sectors is related to

uplift of deep seated zones of the batholiths along

reactivated faulted boundaries. Transitions from granitegreenstone terrains into gneiss—granulite suites involve

a decrease in the abundance of supracrustal enclaves and

an increased strain rate. Whereas early greenstone

sequences are invariably intruded by tonalitic/trondhjemitic/granodioritic gneisses, stratigraphically higher

successions may locally overlap older gneiss terrains and

their entrained xenoliths unconformably. The contiguity
45

ORIGINAL PAGE IS

OF POOR QUALITY

of xenolith patterns suggests their derivation, as relics

of regional mafic-ultramafic volcanic crustal units and

places limits on horizontal movements between individual

crustal blocks.

Fig. 1 - A geological sketch map of the Holenarsipur

greenstone belt, Karnataka (after Naqvi, 1981f J« Geol,

Soc. India, 22:^58-469)
46

Alternative models of granite-greenstone relations are

portrayed in Fig. 2. Model 1 applies to late greenstone

belts overlapping sial whereas model 2 to early belts or

stratigraphically basal volcanic units believed to be

derived from simatic crust. Major detachments along

gneiss-greenstone boundaries and local overfolding and

thrusting suggest horizontal tectonic translations.

These are overprinted by the dominantly vertical tectonic

movements related to the diapriric (magmatic and postmagmatic) uprise of the tonalite/trondhjemite plutons.

The contiguous temporal-spatial grid outlined by the

xenolith swarms constrains major lateral movements of

individual blocks relative to each other, placing limits

on plate tectonics interpretations.

MODEL 1

PK LG . UG LS TGX

dz

MA

Pig. 2 - Alternative

models of gneissgreenstone relationships.

a - model 1 - gneissgreenstone basementcover relations, involving deformed unconformities (du).

b - model 2 - gneissgreenstone relations

involving primary and

deformed intrusive

contacts.

c - model 2 portrayed

in block diagram,

showing transition

from granite-greenstone to gneiss-granulite terrain with

crustal depth.

LG - lower greenstone

UG - upper greenstone

A — acid volcanics &

sediments; TGX - Nagneiss with xenoliths;

PK - late granites;

dz - deformed zone;

LS — late sediments;

0 - orthogneiss;

MA - mafic and anorthositic inclusions;

xb — xenoliths
N86-23106

47

DISMEMBERED ARCHEAN OPHIOLITE IN THE SOUTHEASTERN WIND RIVER

MOUNTAINS, WYOMING — REMAINS OF ARCHEAN OCEANIC CRUST. G.D. Harper,

Dept. Geological Sciences, State University of New York, Albany, NY 12222

Archean mafic and ultramafic rocks occur in the southeastern Wind

River Mountains near Atlantic City, Wyoming (Figure) and are interpreted

to represent a dismembered ophiolite suite. The ophiolitic rocks occur

in a thin belt intruded by the 2.6 Ga Louis Lake Batholith on the

northwest (1, 2). On the southeast they are in fault contact with the

Miners Delight Formation comprised primarily of metagraywackes with minor

calc-alkaline volcanics.

The ophiolitic and associated metasedimentary rocks (Goldman Meadows

Formation) have been multiply deformed and metamorphosed. The most

prominant structures are a pronounced steeply plunging stretching

lineation and steeply dipping foliation. Pillow lavas are stretched

parallel to the lineation and typically have aspect ratios of 10:3:1.

Bedding in banded iron formation shows polyphase folding with fold axes

parallel to the stretching lineation; sheath folds are locally well

developed. The intrusive contact of the Louis Lake batholith with the

ophiolitic rocks has been extensively modified by deformation; the

batholith becomes progessively more deformed as the contact is

approached, and at the contact the batholith is strongly lineated and

mylonitic. The contact between the ophiolitic rocks and the Miners

Delight Formation is a major fault zone (Roundtop Fault) containing

amphibolite-facies mylonites overprinted by greenschist-facies brittle

cataclasites (3). These structural data indicate that the ophiolitic and

associated metasedimentary rocks have been deformed by simple shear when

the Miners Delight was emplaced over the Louis Lake batholith and its

ophiolitic wall rocks.

The ophiolitic rocks include ultramafics, metagabbros, metadiabases,

and pillow lavas. Relict structures and textures are often well

preserved. However, an ophiolite "stratigraphy" is not present; the

ophiolitic rocks consist of tectonic slices, from northwest to southeast,

of (1) metadiabase, (2) metagabbro and ultramafics, (3) pelitic schists,

quartzite, and banded iron formation (Goldman Meadows Formation), and (4)

greenschist and amphibolite (Roundtop Mountain Greenstone) locally

containing pillows and massive flows or sills. In addition, a thin

sliver of pillow lavas occurs between the metadiabase and ultramafic

rocks at one locality, but is separated from the metadiabase by a

strongly foliated talc-actinolite-chlorite schist.

The ultramafic rocks are largely serpentinites, but some have

amphibole-chlorite assemblages and one clinopyroxenite was found. Many

of the ultramafic rocks and associated metagabbros have well-preserved

relict cumulus textures, and igneous layering is visible in a few

outcrops. The ultramafic rocks and associated metagabbros are only

weakly deformed, in contrast to the highly deformed mafic and

metasedimentary rocks.

Metadiabase occurs in a wide belt along the margin of the Louis Lake

batholith, and much of it occurs as large xenoliths within the margin of

the batholith. The metadiabase unit locally contains numerous parallel

dikes, some of which show one-way chilling. Medium to coarse-grained
ARCHEAN OPHIOLITE

48

 G.D. Harper

metagabbro occurs locally within the metadiabase; some of the inetagabbro

occurs as thin screens between fine-grained metadiabase dikes. These

features suggest that the metadiabase unit represents a deformed sheeted

dike complex.

The Roundtop Mountain Greenstone contains common pillow structures

with well-preserved chilled rims. Massive lavas or sills comprise a

significant portion of the formation, and gray phyllites occur rarely.

Rare isolated outcrops of black and foliated "basaltic komatiites,"

consisting primarily of actinolite and chlorite, occur in both the

Roundtop Mountain Greenstone and metadiabase. However, they are

chemically very different from the pillow lavas and metadiabases and

possibly represent younger alkaline dikes.

Metasedimentary rocks of the Goldman Meadows Formation over]ying(?)

the Roundtop Mountain Greenstone consist of pelitic schist, quartzite,

and banded iron formation (1). The banded iron formation possibly formed

by precipitation from hydrothermal vents in a manner similar to modern

metalliferous sediments formed at spreading centers (4). Mafic sills and

dikes (amphibolites) intrude the metasedimentary rocks, and are

themselves deformed and metamorphosed.

Geochemical analyses were made of the metadiabase and pillow lavas to

determine whether they are genetically related (5). "Immobile" trace

element compositions (Ti, V, Cr, Ni, Zr, Y, Nb) are very similar in both

units, consistent with the interpretation that they comprise different

parts of a dismembered ophiolite. These rocks are similar to modern

enriched mid-ocean ridge basalts.

The ophiolitic rocks are interpreted as the remains of Archean

oceanic crust, probably formed at either a mid-ocean ridge or back-arc

basin. All the units of a complete ophiolite are present except for

upper mantle peridotities. The absence of upper mantle rocks may be the

result of detatchment within the crust, rather than within the upper

mantle, during emplacement. This could have been the result of a steeper

geothermal gradient in the Archean oceanic lithosphere, or may have

resulted from a thicker oceanic crust in the Archean (6).

REFERENCES

(1) Bayley, R. W. , Proctor, P. D., and Condie, K. C. (1973). U.S. Geol.

Surv. Prof. Paper 793, 39 pp.

(2) Stuckless, J.S., Hedge, C.E., Worl, R.G., Simmons, K.R., Nkomo, I.T.,

and Wenner, D.G. (1985). Geol. Soc. Am. Bull., 96, p. 850-860.

(3) Hull, J. (1984). Geol. Soc. Am. Abs. with Prog., 16, p. 600.

(4) Bonatti, E. (1981). In: Emiliani, C. (Ed.), The Oceanic Lithosphere, New York, Wiley and Sons, p. 639-686.

(5) Harper, G.D. (1985). Ofioliti (in press).

(6) Sleep, N. H., and Windley, B. F. (1982). Jour. Geol.. 90, p. 363-

379.
ARCHEAN OPHIOLITE

G.D. Harper

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N86-23107

OU . * • • A

PRELIMINARY REPORT ON THE GEOLOGY AND GOLD MINERALIZATION OF THE

SOUTH PASS GRANITE-GREENSTONE TERRAIN, WIND RIVER MOUNTAINS, WESTERN WYOMING

(USA); W.D. Hausel, Geological Survey of Wyoming, Laramle, Wyoming 82071

The South Pass granite-greenstone terrain lies near the southern tip of

the Wind River Mountains of western Wyoming. This Archean supracrustal pile

has been Wyoming's most prolific source of gold and iron ore* From 1962 to

1983, more than 90 million tons of iron ore were recovered from oxide-fades

banded iron formation, and an estimated 325,000 ounces of gold were mined

from metagreywacke-hosted shears and associated placers (!)•

Precambrian rocks at South Pass are unconformably overlain by Paleozoic

sediments along the northeast flank, and a Tertiary pediment buries Archean

supracrustals on the west and south. To the northwest, the supracrustals

terminate against granodiorite of the Louis Lake batholith; to the east, the

supracrustals terminate against granite of the Granite Mountains batholith.

The Louis Lake granodiorite is approximately 2,630 + 20 m.y. old (2), and the

Granite Mountains granite averages 2,600 m.y. old (3).

The geometry of the greenstone belt is best expressed as a synform that

has been modified by complex faulting and folding. Metamorphism is amphibolite grade surrounding a small island of greenschist fades rocks.

The youngest of the Archean supracrustal successions is the Miners

Delight Formation. This unit yielded a Rb-Sr isochron of 2,800 m.y. (2). A

sample of galena from the Snowbird Mine within the Mienrs Delight Formation

yielded a model age averaging 2,750 m.y. (4). The Snowbird mineralization

appears to be syngenetic and is hosted by metavolcanics of calc-alkaline

affinity.

Based on regional mapping by Bayley and others (5) and by the author (in

progress), four mappable supracrustal units are present. The uppermost unit,

the Miners Delight Formation is greater than 1,600 m thick and consists of

metagreywacke, metavolcanics, metaconglomerate, graphitic schist, and

tremolite-actinolite schist. Underlying, and in fault contact with turbidites in the Miners Delight Formation, are metatholeiites of the Roundtop

Mountain Formation. These metatholeiites are amphibolites, greenstones, and

pillow metabasalts. The geometry of the pillows, which has been used for

determining the tops and bottoms of units (5, 6) has only produced ambiguous

conclusions due to the intense deformation.

The Roundtop Mountain greenstones are underlain(?) by quartzite, metapelite, and banded iron formation of the Goldman Meadows Formation. This unit,

in turn, is underlain(?) by mafic and ultramafic schists tentatively named

the Diamond Springs ultramafics. This ultramafic unit consists of amphibolite, serpentinite, metaperidotite, and tremolite-talc-chlorite schist.

Harper (6) interprets this unit to represent a dismembered ophiolite

sequence.

Mining districts occur on both limbs of the South Pass synform. While

the South Pass - Atlantic City District occurs along the northwestern limb,

the Lewiston District is found on the eastern limb (7). Gold mineralization

in the South Pass - Atlantic City District is found chiefly in shear zones in
SOUTH PASS, WYOMING

. , .-'- ./. '

Hausel,''W.D.

 51

metagreywacke adjacent to metagabbro sills and dikes. Wall-rock studies of

the auriferous shears, show Si and K have been enriched and Ca and Mg have

been leached. Mineralogically, these chemical changes are expressed as weak

phyllic alteration of the wall rock. Analyses for native gold from the

Diana Mine show high Au/Ag and low Au/Cu ratios (8). The gold analyses and

wall-rock alteration are characteristic of a hypothermal vein.

The Lewiston District on the eastern flank of the synform includes

strike-trending, metagreywacke - hosted, auriferous shears along the limb of

a major fold (9). A few major lodes are localized where the strike shears

intersect cross-cutting shears. Wall rocks show distinct chloritic and hematitic alteration as well as weak phyllic alteration.

(1) Hausel, W.D., 1980, Gold districts of Wyoming: Geological Survey of

Wyoming Report of Investigations 23, 71 p.

(2) Stuckless, J.S., Hedge, C.E., Worl, R.G., Simmons, K.R., Nkomo, I.T., and

Wenner, D.B., 1985, Isotopic studies of the late Archean plutonic rocks

of the Wind River Range, Wyoming: Geological Society of America Bulletin, v. 96, p. 850-860.

(3) Stuckless, J.S., and Peterman, Z.E., 1977, A summary of the geology,

geochronology, and geochemistry of Archean rocks of the Granite Mountains, Wyoming: Wyoming Geological Association Earth Science Bulletin,

v. 10, no. 3, p. 3-10.

(4) Cannon, R.S., Jr., Bayley, R.W., Stern, T.W., and Pierce, A.P., 1966,

Ancient rocks and ores in south-central Wyoming [abs.]: Geological

Society of America Rocky Mountain Section 18th Annual Meeting Program, p.

27.

(5) Bayley, R.W., Proctor, P.O., and Condie, K.C., 1973, Geology of the South

Pass area, Fremont County, Wyoming: U.S. Geological Survey Professional

Paper 793, 39 p.

(6) Harper, G.D., in press, Dismembered Archean ophiolite, Wind River Mountains, Wyoming (USA), in Desmons, J., ed., Ophiolites through time:

Ophioliti.

(7) Harris, R.E., Hausel, W.D., and Meyer, J.E., 1985, Metallic and

industrial minerals map of Wyoming: Geological Survey of Wyoming Map

Series MS-14, scale 1:500,000.

(8) Antweiler, J.C., and Campbell, W.L., 1977, Application of gold compositional analyses to mineral exploration in the United States: Journal

of Geochemical Exploration v. 8, p. 17-29.

(9) Hausel, W.D., 1984, Preliminary geologic map of the Lewiston gold mining

district (Radium Springs Quadrangle), South Pass, Wyoming: Geological

Survey of Wyoming unpublished mineral report # MR 84-7, scale 1:24,000.
N86-23108

THE KOLAR SCHIST BELT: A POSSIBLE ARCHEAN SUTURE ZONE

G. N. Hanson1, E. J. Krogstad1, V. Rajamanl2 and S. Bal akrlshnan2, (1) Department of Earth and Space Sciences, SUNY, Stony Brook, NY 11794 (2) School of

Environmental Sciences, Jawaharlal Nehru University, New Delhi 110067, India.

The Kol ar Schist Belt 1n the Karnataka craton, south India, Is a 4 to 20

km by 80 km long, N-S trending Archean supracrustal belt dominated by mafic

metavolcanlcs. The schist belt 1s surrounded on both sides by granodlorltlc

gneisses collectively known as the Peninsular Gneiss. Our work has shown that

the Kolar Schist Belt and the surrounding gneisses Include major

discontinuities 1n age, structural style, and composition. These

discontinuities are defined by the schist belt Itself.

The results reported here are based on our Rb-Sr, Sm-Nd, and Pb-Pb whole

rock Isotope data; U-Pb dating of zircon and sphene; major and trace element

(Including REE) analyses; and field observations.

The schist belt 1s broadly synformal, but 1s complexly refolded Into

basin and dome structures (D. Mukhopadyay, personal communication). The first

period Involved N-S trending Isoclinal recumbant folds during E-W compression.

These folds were refolded Into tight, upright folds along E-W trending axes.

This sequence 1s broadly similar to those seen 1n other schist be! ts 1n the

western part of the Karnataka craton.

Contacts between the Peninsular Gneiss and the margins of the belt have

long been thought to represent an eroslonal unconformity. However, our recent

field work Indicates that the rocks at the contacts are physically Interleaved

by left lateral shearing. Due to this shearing the adjoining gneisses have

been converted to quartz-muscovlte schists, which were previously Interpreted

to be metasedlmentary rocks.

The gneisses east of the schist belt are relatively homogeneous,

granodlorltlc gneisses which were folded prior to Intrusion of minor fel sic

bodies. Folds have not yet been defined 1n these gneisses, but a strong

foliation was developed which strikes NNE and dips steeply to the west,

suggesting horizontal compression.

The gneisses west of the schist bel t show a much more compl ex, earl 1er

history than that of the eastern gneisses. The granodlorltlc Dod Gneiss 1s the

earliest unit on the western side of the schist belt. This rock was subjected

to a period of deformation shown by an early foliation seen 1n some lessstrained exposures. Subsequently, the Dod Gneiss was Intruded by the

leucocratlc, granodlorltlc Dosa Gneiss and the granodlorltlc Patna Granite.

Fol 1 owing the Intrusion of the Dosa Gneiss, the terrane to the west of

the schist belt was subjected to a period of horizontal compression producing

tight to Isoclinal, W overturned folds with gently N or S plunging axes. The

strong NNE axial planar foliation produced by this deformation 1s cut by the

later N-S shears along the western margin of the schist belt.
THE KOLAR SCHIST Bai

 53

Hanson, G. N., e.t al.

The gneisses on the east and west side of the belt have been dated using

U-Pb ages for small populations of abraded zircons and abraded single zircons

as well as sphene. These zircons commonly give concordant ages. In which case

the small populations of zircons (ca. 100 mlcrograms) have analytical

uncertainties of less than 1 Ma. and the single zircons have uncertainties of

about 5 Ma.

Gneisses east of the belt were Intruded at 252911 Ma based on U-Pb ages

for zircon. This age 1s consistent with the Rb-Sr and Pb/Pb whole rock

Isochron ages. The Isochrons have a mantle-like Initial ratio for Sr

(87/86=0.7013) and mu=8 for the Pb data. These values suggest that the

gneisses were not derived from a much older continental crust. U-Pb ages for

metamorphlc sphene are 2520 ±1 Ma suggesting that the gneisses were

metamorphosed to at least amphlbol1te grade at that time.

West of the belt, based on U-Pb ages for zircon, the Dod Gneiss was

empl aced at 2610+5 Ma, the Dosa Gneiss was Intruded at 2550±10 Ma and the

Patna Granite at 2551+1 Ma. The time of metamorphlsm based on the U-Pb ages

for sphene from the Dod Gneiss 1s 2551±1 Ma. Rb-Sr and Pb/Pb whole rock data

suggest that the gneisses were variably contaminated by an older basement. UPb ages for some of the single zircon cores from the Dod Gneiss and later

aplltlc dikes Indicate a zircon component was Inherited from this basement,

which has a minimum age of 3200 Ma. The basement, which has not yet been

clearly Identified 1n the field, seems to Include quite evolved felslc rocks.

In the Kolar Schist Belt there are two suites of komat11t1c and

tholelltlc amphlbol 1tes. Both the komat11t1c and thole11t1c amphlbol 1tes on

the eastern side are light REE enriched, and almost all of the komat11t1c and

tholelltlc amphlbol 1tes 1n the west-central part of the belt are lightest REE

depleted. The preservation of rare pillow structures and the association of

the amphlbol1tes with Iron formation suggest that the amphlbol 1tes were formed

under submarine conditions. The grade of metamorphlsm 1s amphlbol 1te fades.

Rajamanl &t .al. (1) concluded that the komat11t1c amphlbol 1tes from both

the east and west central part of the belt were derived by 10 to 25% melting

at depths greater than 80 km and at temperatures greater than 1500°C 1n a

mantle with an FeO/MgO ratio greater than that of pyrol 1te. Other models

proposed for the generation of komatlltes generally require larger percentages

of melting to generate the high MgO abundances.

Rajamanl .e± jj.. (1 and 1n preparation) suggest that the thole11tes appear

to have been derived by melting at shallower levels than the komatlltes and

derived from sources which were highly variable 1n their FeO/MgO ratios,

generally with FeO/MgO ratios much greater than that for the sources for the

komatlltes. The key arguments are that: the tholelites are very Iron-enriched
54

THE KOLAR SCHIST

Hanson, G. N., et al.

compared to the field for potential melts of pyrol 1te at pressures less than

25 kb on an ol1v1ne saturation surface; and while the Incompatible elements

show similar ratios 1n the komatlltes and thol elites for each suite, the

expected correlations between major and trace elements for differentiation

from komatlltes or melting of sources similar to those of komatlltes are not

found.

Sm-Nd data for komatlltes from both sides of the belt 1 1e with large

variations about a 2900 Ma Isochron. It 1s not clear why the data lie about a

2900 Ma Isochron. Is this the age of these amphlbol 1tes? If this 1s so, they

are much older than the Igneous felslc rocks on either side of the belt which

are 2500 to 2600 Ma. Or, 1s this the time when the sources became variably

light REE enriched and depleted? Some of the variation 1n the Sm/Nd ratios 1s

clearly a function of melting processes 1n which garnet was left 1n the

residue. Perhaps the variability 1n the data about the reference line reflects

a number of reasons such as: variable times of light REE depletion and

enrichment of their mantle sources; as well as the possible effects of crustal

contamination or metamorphlc alteration.

Even though the ages of the units making up the Kol ar Schist Belt are

poorly constrained, the sources of the amphlbol1tes so far analyzed had longterm histories of LREE depl etlon (epsl 1 on Nd of +2 to +8 for an age of 2900

m.y.) relatlve to other Archean mafic rocks which commonly have epsll on Nd

equal to about+2.0 ± 2.0.

The Kolar Schist Belt represents a N-S trending discontinuity 1n the

structures, llthologles, and emplacement and metamorphlc ages of late Archean

gneisses. The suggestion of a much older basement on the west side of the

belt 1s not seen on the east Within the schist belt amphlbol 1tes from each

side have distinctly different chemical characteristics, suggesting different

sources at similar mantle depths. These amphlbol 1tes were probably not part

of a single volcanic sequence, but may have formed about the same time 1n two

completely different settings. Could the amphlbol 1tes with depleted light REE

patterns represent Archean ocean floor volcanlcs which are derived from a

mantle source with a long term depletion of the light REE? Why are the

amphlbol1tes giving an age which may be older than the exposed gneisses

Immediately on either side of the belt? These results suggest that 1t 1s

necessary to seriously consider whether the Kolar Schist Belt may be a suture

between two late Archean continental terranes.

REFERENCES

(1) Rajamanl V., Shlvkumar K., Hanson G. N., and Shlrey S. B. (1985)

Geochemistry and petrogenesls of amphlbol 1tes, Kolar Schist Belt, South India:

Evidence for komat11t1c magma derived by low percentages of me! ting of the

mantl e J. Petrol. 26, p. 92-123.
N86-23109 tJ O

EVIDENCE FOR SPREADING IN THE LOWER KAM GROUP OF THE YELLOUKNIFE

GREENSTONE BELT: IMPLICATIONS FOR ARCHEAN BASIN EVOLUTION IN THE SLAVE

PROVINCE. H. Helmstaedt and W.A. Padgham, Dept. of Geological Sciences,

Queen's University, Kingston, Canada K7L 3N6, and Geology Division, Northern

Affairs Program, P.O. Box 1500, Yellowknife, N.U.T., Canada X1A 2R3

The Yellowkni-fe greenstone belt is located in the southwestern part of

the Slave Structural Province, a Late Archean (2.7-2.5 Ga) granitegreenstone terrane in the northwestern part of the Canadian Shield.

Supracrustal rocks withi n this province, collectivel y referred to as

Yellowknife Supergroup (Henderson, 1970), differ from the supracrustal

successions of the Superior Province and other older Archean terranes by the

absence of komatiites and the high proportion of metasedimentary to

metavolcanic rocks. The Yellowknife belt was first mapped by Jolliffe (1942,

1946) on the scale of one inch to one mile , and the gold-producing area

around Yellowknife was remapped on a more detailed scale (1:12 000) by

Henderson and Brown (1966). As the belt became the best-known example of the

basalt-dominated supracrustal belts in the western Slave Province (Padgham,

1985), the stratigraphic framework established here (Henderson, 1970), formed

the basis for the development of models for Archean basin evolution (McGlynn

and Henderson, 1972; Henderson, 1981). Under a r-ecent mapping program of the

Geology Divisio n of the Northern Affairs Department in Yellowknife, detailed

mapping was extended, and a 1:10 000 map series for the entire belt is

currently under preparation. This work resulted in a number of revisions and

refinements in the established stratigraphy (Helmstaedt and Padgham, 1986)

and provides the basis for a reassessment of current models of greenstone

belt evolution in the Slave Province.

The major portion of the Yellowknife greenstone belt is underlain by the

predominantly mafic rocks of the Kam Group which consists of a northeasterlystriking, homoclinal sequence of flows and tuffs that dip steeply and face

uniformly to the southeast (Fig. 1). Numerous dikes, sill s and irregular

bodies of gabbro and locally anorthosite appear to form an integral part of

the volcanic sequence. The Kam Group has been subdivided into four

formations (Fig. 2) with a combined thickness of approximately llkm. The

lower contact is obscured by the intrusion of a composite batholith (Western

Granodiorite, Fig. 1) that cuts across the strike of the flows. At the base

of the exposed section, near the northern end of the belt, a narrow band of

felsic volcanic rocks and banded iron- formation is in conformable contact

with overlying pillowed flows above which a mafic extrusive- intrusive

complex is developed (Fig. 2) whose pseudostratigraphy resembles that of

certain Phanerozoic ophiolites. Near the southwestern end of the belt, the

upper part of the Kam Group (Yellowknife Bay Formation) overlaps a sequence

of older volcanic and sedimentary rocks belonging to the Octopus Formation

(Fig. 1). In the northern part of the belt, the upper formations of the Kam

Group are truncated by an unconformity beneath conglomerates and sandstones

of the Jackson Lake Formation. Farther to the south, where the top of the

Kam is preserved locally, it is overlain by calc-alkaline rocks of the

Banting Group that, in turn, are overlain by turbidites of the Walsh and

Burwash Formations. All rocks of the Yellowknife Supergroup are deformed and

metamorphosed, with metamorphic grade increasing from greenschist to

amphibolite facies towards the granitoid intrusions. Inspite of the

metamorphic overprint, however, primary structures and intrusive

relationships are well preserved.
^ • . . . •- '. • i , s? I

56

The mafic intrusive-extrusive complex of the Chan Formation <Fig. 2)

grades -from a lower part, dominated by gabbro, through a multipl e dike

complex into massive and pillowed flows with thi n beds o-f inter-flow

sediments. At the base o-f the section is a sheet-like body o-f massive,

medium- to coarse-grained, locally layered gabbro that was intruded into a

sequence o-f pillowed -flows, remnants of which are preserved at three levels.

The upper boundary of this body is a relatively sharp transition into the

dike complex which consists of numerous, fine- to medium-grained metadiabase

dikes and septa and irregular bodies of relatively coarse gabbro between

which screens of pillowed flows can be recognized. The dikes, which are

locally sheeted, show symmetric and asymmetric chi 1 led margins and range in

width from less than one to over 10m. Some dikes grade into pillows,

suggesting that they were intruded close to the seafloor and may have acted

as feeder system to the growing volcanic pil e <de Wit and Stern, 1978). Most

of the irregular gabbros are multipl e intrusions with abundant chille d

margins and extremely complex contact relationships. Igneous layering is

generally absent at this level, but an up to 100m thick, sheet-like body of

gabbroic anorthosite was recognized (Fig. 2). It is surrounded entirely by

gabbro that has chille d margins against the anorthosite. Though massive and

pillowed flows predominate above the dike complex, sill s and irregular bodies

of gabbro, many of them multipl e intrusions, are common in the upper parts of

the Chan Formation. The top half of the Kam Group continues to be dominated

by pillowed and massive mafic flows, but contains numerous intercalations of

felsic tuffs and tuffaceous sediments. Some of the flows and many of the

interflow tuffs and sediments are continuous along strike for more than 10 km

and allow stratigraphic correlation across Proterozoic transcurrent faults

(Fig. 1). Synvolcanic mafic intrusions in this part of the section consist

of numerous sill s some of which are connected to dike swarms. The entir e

section was intruded also by several post-volcanic dike swarms.

The Yellowknife greenstone belt has been interpreted as the western

margin of an Archean turbidite-fi1 led basin bordered in the east by the

Cameron River and Beaulieu River volcanic belts (Henderson, 1981; Lambert,

1982). This model implie s that rift ing was entirely ensialic and did not

proceed beyond the graben stage. Volean ism is assumed to have been

restricted to the boundary faults, and the basin was floored by a downfaulted graniti c basement. On the other hand, the enormous thickness of

submarine volcanic rocks and the presence of a spreading complex at the base

of the Kam Group suggest that volcanic rocks were much more widespread than

indicated by their present distribution. Rather than resembling volcanic

sequences in intracratonic graben structures, the Kam Group and its tectonic

setting withi n the Yellowknife greenstone belt have greater affinitie s to the

Rocas Verdes of southern Chil e (deUit and Stern, 1981), Mesozoic ophiolites,

that were formed in an arc-related marginal basin setting. The similaritie s

of these ophiolites with some Archean volcanic sequences was previously

recognized by Tarney et al . (1976) and served as basis for their

marginal-basin model of greenstone belts. The discovery of a multipl e and

sheeted dike complex in the Kam Group confirms that features typical of

Phanerozoic ophiolites are indeed preserved in some greenstone belts and

provides further fiel d evidence in support of such a model.

FIGURE CAPTIONS: (1). Geological map of the Yellowknife greenstone belt.

Modified from published maps of the Geological Survey of Canada and Northern

Affairs Program, Yellowknife. (2). Generalized section of the Kam Group.
57

REFERENCES:

de Wit , M.J. and Stern, C.R. (1978) Jour. Volcanology and Geothermal Res., 4,

p. 55-80.

de Wit , M.J. and Stern, C.R. (1981) Tectonophysics, 72, p. 229-260.

Helmstaedt, H. and Padgham, W.A. (1986) Canadian Jour. Earth Sciences (in

press).

Henderson, J.B. (1970) Geol. Surv. Canada, Paper 70-26.

Henderson, J.B. (1981) in Kroner, A., ed., Plate Tectonics in the

Precambrian. Elsevier, Amsterdam, p. 213-235.

Henderson, J.F. and Brown, I.C. (1966) Geol. Surv. Canada, Bull. 141.

Jolli-He, A.UI. (1942) Geol. Surv. Canada, Map 868A.

Jolli-ffe, A.U. (1946) Geol. Surv. Canada, Map 709A.

Lambert, M.B. (1982) Geol. Surv. Canada, Paper 82-1A, p. 165-167.

McGlynn, J.C. and Henderson, J.B. (1972) in Geol. Assoc. Canada, Special

Paper 11, p. 506-526.

Padgham, W.A. (1985) in Geol. Assoc. Canada, Special Paper 28, p. 133-151.

Tarney, J., Dalziel, I.W.D., and de Wit , M.J. (1976) jm Uindley, B.F., ed.,

The Early History oi the Earth. Uliley, London, p. 131-146.
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N86-23HO 59

ARCHAEAN WRENCH-FAULT TECTONICS IN THE ABITIBI

GREENSTONE BELT OF CANADA; C. Hubert and J.N. Ludden. Department

of Geology, Universite de Montreal, Montreal, Canada, H3C 3J7.

Structural studies in the southern sector of the Abitibi

greenstone belt of Canada have defined a deformation style

associated with a wrench-fault system (1 ). The fundamental

features of this tectonic regime are the following:

i) the formation of lozenge - shaped blocks of terrane

which are bounded either by fault - zones or by highly strained

zones of ductile deformation. In these blocks there is a

pronounced gradient in degree of deformation from well preserved

cores to highly deformed and sometimes mylonitized margins;

ii) sedimentary accumulations occur along the margins

of the blocks in a series of narrow basins bounded by shear -

zones;

iii) blocks of different lithologies and structural and

metamorphic histories have been juxtaposed;

The deformation history is summarized below and shown in

simplified form in Figure 1. The first deformation phase was

simple shearing associated with WSW - ESE sinistral wrench

faulting which resulted in NW - SE fold traces and transected

schitosities. Progressive deformation affected blocks of terrane

in a tectonic regime in which volcanism, shearing, deformation

and uplift and erosion were synchronous; terranes composed

dominantly of felsic volcanics were juxtaposed with blocks of

ultramafic volcanic and sedimentary accumulations.

The first deformation phase was followed by N - S

compression resulting in the development of major E - W thrustshears. This deformation resulted in the formation of an E - W

fold trace and crenulation clevage. The superposition of the two

deformation episodes resulted in the generation of NE - SW and NW

- SE complementary faults defining " S " and " Z " sigmoidal

forms in highly strained E - W shear - zones ( Figure 2).

U - Pb zircon ages, compiled in Ludden et al., ( 2 )

indicate that the volcanic accumulations in the Porcupine, RouynNoranda and Val D1 Or areas of the southern Abitibi belt define

an axis of volcanism of tholeiitic lineage that was at its peak

at approximately 2700 m.y.. These volcanic rocks superimpose an

older volcano-plutonic terrane which is characterized in the NE -

Abitibi belt and can be correlated towards the SW across the

Kapuskasing front to the Wawa subprovince ( 2,3 ). This axis of

volcanism is approximately 2850 - 2720 m.y. in age and is

dominated by calcalkaline volcanic and plutonic rocks.
60

ARCHAEAN WRENCH-FAULT TECTONICS

C. Hubert and J.N. Ludden

A tectonic model is proposed in which the southern Abitibi

belt formed in a series of rift basins which dissected an

earlier formed volcanic arc. Comparisons can be made with

Phanerozoic areas such as, the Hokuroko basin of Japan, the Taupo

volcanic zone of New Zealand and the Sumatra and Nicaragua

volcanic arcs. In addition the identification of the major E - W

thrust shears make it possible to speculate that the southern

Abitibi belt comprises a collage of blocks of terrane which have

been accreted against a more stable continental margin or microcontinent. If this interpretation is correct analogies can be

made with the SW margin of the U.S.A. in which recently formed

blocks of volcanic terrane are being accreted against the

western margin of the U.S.A..

FIGURE 1 : Deformation History of the Southern Abitibi Belt.

(BLOCK-I)

WNW-ESE SINISTRAL

SHEAR ZONE

(BLOCK 2)

E-W THRUST SHEAR/

(BLOCKS)

SEDIMENTS
c

ARCHAEAN WRENCH-FAULT TECTONICS

C. Hubert and J.N. Ludden

ORIGINAL PAGE IS

OF POOR QUALBTY

FIGURE 2 : Schematic representation of " lozenge - shaped "

blocks of terrane bounded by shear-zones and thrust-shears In the

Southern Abitlbl belt

49'

00

79°00 78°00

48'

00

!0v VFELSIC

loj y COMPLEX

j°

 V

•-w^X O o O^-^C^^iSs-O

~~^HT~ r ° o ° v ..ft**. ,~~~ MAFIC

^>o

FELSICj

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yisiili-H&s''''1'**- ^^ff^:'

:

!^'

y^•vA^^v;•'•.V.':'.•^••V:••:•.:.•':•.'••.'•::'••.•.::;

49°

00

79°00

INTRUSIVE ROCKS

|;X;'f ] Q!MONZONITE-GRANOOIORITE

|^'***^ TONALITE-TRONDJHEMITE

GNEISS

78°00

50km

VOLCANIC ROCKS

|V y| FELSIC VOLCANIC PREDOMINANT

|O p| MAFIC VOLCANIC PREDOMINANT

ULTRAMAFIC

SEDIMENTARY ROCKS

[ H LATE SEDIMENTARY

1

 SERIES

SEDIMENTARY WEDGE

SYMBOLS

61

SHEAR ZONE

REFERENCES:

(1) Hubert C., Gelinas L., Trudel P., 1984, Can J. Earth Sci..

22. 240-255.

(2) Ludden J.N., Hubert, C., Garlepy, C., 1986 , Geol. Mag., In

Press.

(3) Perclval J.A. and Krogh T.E. , 1983, Can J. Earth Sci., 20.

830-843.
N86-23111

62

TRANSPRESSION AS THE MAIN DEFORMATIONAL EVENT IN AN ARCHEAN

GREENSTONE BELT, NORTHEASTERN MINNESOTA; P.J. Hudleston and D. SchultzEla, Department of Geology and Geophysics, University of Minnesota,

Minneapolis, MN 55455; R.L. Bauer, Department of Geology, University

of Missouri, Columbia, MO 65211; D.L. Southwick, Minnesota Geological

Survey, 2642 University Ave., St. Paul, MN 55114.

Deformed and metamorphosed sedimentary and volcanic rocks of the

Vermilion district constitute an Archean greenstone belt trending eastwest between higher grade rocks of the Vermilion Granitic Complex to

the north and the Giants Range batholith to the south. Metamorphic grade

is low throughout, being lowest in the center of the belt (chlorite

zone of the greenschist facies) (1). All the measured strain, a cleavage

or schistosity, and a mineral lineation in this belt are attributed

to the 'main1 phase of deformation (Dp) (2) that followed an earlier

nappe-forming event (D1)(3, 4), which left little evidence of penetrative

fabric (2).

 i

Previous work assumed that the D~ deformation resulted from northsouth compression across the district, presumably, related to diapiric

intrusion of the batholithic bodies to the north and south (1). A number

of lines of evidence now lead us to believe that a significant component

of this deformation resulted from dextral shear across the whole region.

Thus the Vermilion fault, a late-stage largely strike-slip structure

(1) that bounds the Vermilion district to the north, may simply be the

latest, most brittle expression of a shear regime that was much more

widespread in space and time. Features that are indicative of shear

include ductile shear zones with sigmoidal foliation patterns, highly

schistose zones with the development of shear bands, feldspar clasts

or pyrite cubes with asymmetric pressure shadows, and the fact that

the asymmetry of the F~ folds is predominantly Z for at least 15 km

south of the Vermilion fault.

The presence of a large component of simple shear may help explain

additional structural features in a simpler way than otherwise possible.

Just south of the Vermilion fault the cleavage locally becomes folded

and a new spaced cleavage develops in a similar orientation to the old

cleavage away from the folds. Rather than interpreting this as evidence

for an additional episode of deformation, we consider it to be due to

a single process of continuous shear: a foliation develops and after

a large strain local perturbations result in folding of the old foliation

and the development of a new one axial planar to the folds.

The same type of perturbation can lead to the juxtaposition of

ENE-trending zones of constrictional and flattening strains (5), a distinctive feature of the rocks of the Vermilion district otherwise hard

to account for. The maximum extension directions (X) of all samples

showing constrictional strain, plunge east at angles between 30 and

65°. X in samples showing flattening strain plunges east or west, but

near the Vermilion fault all plunges are west or more steeply east than

they are in constrictional samples. The maximum shortening direction

(Z) plunges consistently less than 25° to the north or south.
1TRANSPRESSION, NE MINNESOTA

Hudleston, P.J. et al.

 63

The strain variations require a model which can satisfy compatibility

constraints and space considerations. The area of consistent constrictional strains in the south may represent one regional component of

the strain. Spatial correspondence of flattening strains with the Vermilion fault suggests that a simple shear component was added in that

area. A modified model of transpression may explain how E-plunging X

axes are reoriented to become W-plunging by a concomitant inhomogeneous

progressive simple shear. Less than vertical plunge of the X axes may

necessitate some component of oblique motion on the fault.

In a general way the strain patterns observed in the Vermilion

'district can be reasonably explained by a history of N-S shortening

accompanied by inhomogeneous dextral simple shear. The variations of

strain may be a consequence of variations in the relative intensities

of shortening and shear, large perturbations of the shear, or the influences of other structures. There may be an analogy with the strain

partitioning that occurs in small scale ductile shear zones at large

strains.

For transpression to have occurred, the Vermilion district would

have to have been a region of relatively soft crust caught beween two

more rigid (either thicker or cooler) blocks to north and south. We

do not yet know to what extent the high-grade terranes to north and

south were also affected by transpression deformation and therefore

the configuration of the more rigid block.

References

(1) Sims, P.K. (1976) Bull. Geol. Soc. Am.. 87, p. 379-389.

(2) Hudleston, P.J. (1976) Can. J. Earth Sci.. 13_, p. 579-592.

(3) Bauer, R.L. (1985) Geology, 13, p. 657-660.

(4) Bauer, R.L., Hudleston P.J. and Southwick d.L. (1986) In Workshop

on the Tectonic Evolution of Greenstone Belts. Lunar and Planetary

Institute, Houston (this volume).

(5) Ela, D. and Hudleston P.J. (1985) In 31st Institute on Lake Superior

Geology (abstract).
N86-23112

64

A CONTINUOUS RECORD OF TECTONIC EVOLUTION FROM 3.5 Ga

TO 2.6 Ga IN SWAZILAND AND NORTHERN NATAL

D.R. Hunter, A.H. Wilson, J.A. Versfeld, A.R. Allen, R.G. Smith,

D.W.W. Sleigh, P.B. Groenewald, G.M. Chutter and V.A. Preston,

Department of Geology, University of Natal (Pietermaritzburg).

The 'vS.S Ga-old bimodal suite underlying an extensive area in southwestern Swaziland comprises the oldest-dated sialic rocks in the Kaapvaal

structural provincep). The suite consists of leucocratic, layered

tonalitic-trondhjemitic gneisses and amphibolites characterized by the

effects of repeated high strains(2). This suite is considered to represent

a sialic basement on which metavolcanic and metasedimentary rocks, now

preserved as scattered 'greenstone1 remnants, accumulated. Direct evidence

to confirm this temporal relationship is lacking, but structural data from

the Dwalile, Assegaai and Commondale areas indicate that (i) the bimodal

gneisses experienced a complex structural history prior to the first

recognizable deformation in the supracrustal rocks (i.e. D1 in the

supracrustals is equivalent to Dn + 1 in the gneisses) and (ii) scattered

remnants of the Dwalile rocks infolded with the bimodal suite structurally

overlie the gneisses and are preserved in synformal keels (2)(3).

Significant proportions of metaquartzites and metapelites are present in

the Assegaai 'greenstone' sequence, the presence of which implies the

existence of felsic crust in the source area from which these sediments

were derived, a conclusion that is consistent with the structural data.

Ultramafic and pillowed mafic rocks of komatiitic and

tholeiitic affinity are present in all four 'greenstone' remnants, but each

contains distinctive lithologies. The Assegaai sequence is characterized

by the abundance of clastic and chemical sediments that are a minor

component of the Commondale and Nondweni remnants. In the former there is

a prominent sub-volcanic intrusion composed of multiple layers of massive

serpentinite (in which relict cumulate olivine is present locally)

alternating with spinifex-textured (olivine and pyroxene) layers. There is

a consistent relationship in the thicknesses of the individual layers, i.e.

where the serpentinite layers range from 10 to 40 m in thickness the

spinifex-textured layers are 1 to 3 m thick. At Nondweni the sequence is

dominated by pillowed tholeiites interlayered with high-magnesium basalts

and basaltic komatiites (up to 22% MgO). The latter show well developed

pyroxene spinifex but peridotite komatiites and units with olivine spinifex

are entirely absent. Silicification of the volcanics considered to be

contemporaneous with extrusion is not uncommon. Within the volcanic

sequence are numerous graded air-fall tuffs and flows of rhyolite

compositions. A zone with biogenic or stromatolitic structures is also

preserved.

These subtle lithologic differences may reflect different levels of

exposure and/or ages of accumulation. The Nondweni greenstones show a

consistent northwesterly younging direction in rocks which are not highly

strained and which are separated by poorly exposed areas of high strain,

suggestive of tectonic interslicing. In contrast the Assegaai and

Commondale rocks show evidence of early reclined folds, which may be a

reflection of deeper infolding. Preliminary geochronologic data indicate
TECTONIC EVOLUTION IN SWAZILAND AND NORTHERN NATAL

D.R. HUNTER, et al.

 65

that the Dwalile 'greenstones' are of similar age to the Barberton

greenstones(1). Pb-Pb isotopic data from a single komatiitic flow at

Nondweni define an age of 3.15 Ga that is consistent with an Rb-Sr age of

•^3.1 Ga for an associated rhyolite(4). However, Sm-Nd data define an age

of 3.6 Ga for komatiitic, tholeiitic and rhyolitic flows. Possible

explanations are that either the Pb-Pb and Rb-Sr systems were reset at

'v-S.I Ga subsequent to extrusion at %3.6 Ga, or, on eruption 3.1 Ga ago,

the extrusions interacted with -v 3.5 Ga-old felsic crust leading to a range

of initial Nd isotopic compositions of the mafic rocks and the generation

of rhyolites by remelting of that crust(4).

Subsequent to the D1 event (Table 1), mantle-derived tonal itic plutons

(Tsawela and Braunschweig) and the meta-anorthositic Mponono layered

intrusive sheet were emplaced into the bimodal gneisses and Dwalile

greenstones. All these rocks were strongly and repeatedly deformed under

amphibolite-facies conditions (Table 1).

Sheet-like granitoid batholiths were intruded at -\,3.2 and ^3.0 Ga,

the locus of emplacement migrating northwards with decreasing age. The

^3.2 Ga-old multiphase sodic granitoid intrusion screens the Assegaai and

Commondale greenstone remnants from their underlying gneissic basement.

Intrusion occurred in the interval between D1 and 02 in the Assegaai and

Commondale areas. A chemically and mineralogically similar granite also

intrudes the Nondweni 'greenstones' but neither its age nor structural

style have yet been studied.

At a high structural level, a second sheet-like, but more potassic

granite, the vast multiphase Lochiel batholith, was intruded at ^3.0 Ga

north of Dwalile. Following this period of widespread emplacement of

granitic magmas emergence above sea-level of stable continental crust took

place. Subaerial weathering of this dominantly granitoid terrane was

accompanied in the north by the development of braided stream systems

draining southeast off the flank of the NE-trending Lochiel batholith(5)

into the Pongola basin(6). Minor contemporaneous volcanism accompanied the

fluvial sedimentation and heralded a period of subaerial extrusion of lavas

(the 2.94 Ga-old Nsuze Group), that range in composition from basalt to

rhyolite and attain a thickness of ^ 8.5 km SE of Piet Retief(7). No

ultramafic nor high-MgO flow units are present and the sequence is

characterized by the simultaneous extrusion of mafic and acidic lavas.

Typically porphyritic andesites are also present.

The Nsuze Group is preserved in a series of inliers in the south where

its thickness decreases in part due to truncation by the upper (Mozaan)

group of the Pongola Supergroup or by the Palaeozoic Natal Group. Volcanic

rocks are less abundant in the southern inliers. Shallow water subtidal

and tidal-flat sediments including stromatolitic carbonate sands are

prominent in the Wit Mfolozi inlier. A heterolithic unit 1.5 km thick

dominated by pyroclastic rocks interlayered with shallow marine sediments

forms the base of the Nsuze Group south of Babanango. This unit is

truncated towards the east by a 4.0 km thick sequence of tidalite

sediments with interlayers of basaltic andesite lavas. Transport

directions in the inliers are from the north and northwest.
TECTONIC EVOLUTION IN SWAZILAND AND NORTHERN NATAL

66 D.R. HUNTEiR, et al.

Sedimentation in the Mozaan group was largely controlled by the

interaction of a braided alluvial plain and a rnacrotidal basin(8). Mozaan

sediments are not preserved south of the Wit Mfolozi inlier either as a

result of removal by erosion or of non-deposition.

The Mozaan Group is typically deformed into gently dipping, doubly

plunging synclinal structures resulting from interference of NW and NEtrending axial traces. Adjacent to the southern margin of the Kaapvaal

Province, tight E-trending folds with vertical axial surfaces are dominant

reflecting a response to deformation related to the development of the

Natal thrust zone at ^1.1 Ga. The Nsuze Group is highly strained adjacent

to the Swaziland border apparently related to a 20 km wide belt of NWtrending folds and faults with left-lateral movement within which the dykelike, mafic Usushwana Intrusive Suite was emplaced at ^,2.87 Ga(9).

The significance of the Pongola Supergroup lies in the fact that it

demonstrates the co-existence of stable continental crust in southeastern

Africa and metastable crustal conditions in southern central Africa

dominated by extrusion and intrusion of voluminous komatiitic and

tholeiitic magmas.

Emplacement of large volumes of granitic magmas principally into

Pongola rocks terminated Archaean evolution. Multiple gneiss domes

separated by screens of Mozaan sediments of high metamorphic grade

developed in southern Swaziland adjacent to the belt of NW-striking, highly

strained Nsuze rocks. Subsequently a thin sheet (300 to 1000 m thick) of

potassic granite was emplaced at the unconformity between the Mozaan Group

and its gneissic granitoid basement. The final pulses of granite plutonism

resulted in the emplacement of sharply transgressive, typically coarsegrained, porphyritic plutons ranging in size from 40 km2 to 650 km2 about

which narrow contact aureoles are developed in the Mozaan sediments. Rb-Sr

isotopic data have yielded only whole-rock errorchrons for these rocks(10).

The concentration of post-Pongola granitoids within the core of the

Pongola depository suggests that depression of the depositional basin

promoted partial melting of the lower crust, which would be consistent with

the proposed model for the genesis of the granitic melts based on

geochemical data(11). The post-Pongola granites differ in their setting

from other Archaean granites in southern Africa(5).

1. Carlson R.W. Hunter D.R. and Barker F. (1985) Abstracts, Spring

Meeting, Amer. Geophys. Union.

2. Jackson M.P.A. (1984). In Precambrian Tectonics Illustrated

E. Schweizerbart'sche Verlagsbachhandlung, Stuttgart, p 1-18.

3. Talbot C.J. Hunter D.R. and Allen A.R. (in press) Deformation of the

Assegaai supracrustals and adjoining granitoid rocks.

4. Wilson A.H. Carlson R.W. Versfeld J.A. and Hunter D.R. (1985).

Abstracts. Fall Meeting, Amer. Geophys. Union.

5. Hunter D.R. (1974) Precambrian Res.. 1, p 259-326.

6. Watchorn M.B. and Armstrong N.V. (1980). Trans. Geol. Soc. S. Afr.,

83, p 231-238.

7. Armstrong N.V. Hunter D.R. and Wilson A.H. (1982). Precambrian Res.,

19, p 75-107.
TECTONIC EVOLUTION IN SWAZILAND AND NORTHERN NATAL 67

D.R. HUNTER, et al.

8. Watchorn M.B. (1980) Precambrian Res.. 13, p 27-42.

9.' Hegner E. Kroner A and Hoffman A.W. (iy84). Earth Planet.-Sci. Lett..

70, p 267-279

10. Barton Jr. J.M. Hunter D.R. Jackson M.P.A. and Wilson A.C.

Trans. Geol. Soc. S. Afr., 86, p 71-80.

11. Condie K.C. and Hunter D.R. (1976). Earth Planet. Sci. Lett., 29, p

389-400.
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N86-23113

IS THE CAMERON RIVER GREENSTONE BELT ALLOCHTHONOUS?;

Timothy M. K,usky, Department of Earth and Planetary Sciences,

The Johns Hopkins University, Baltimore, Maryland, 21218, U.S.A.

Many tectonic models for the Slave Province, N.W.T., Canada,

and for Archean granite - greenstone terranes in general, are

implicitly dependent on the assumption that greenstone belt

lithologies rest unconformably upon older gneissic basement.

Other models require originally large separations between

gneissic terranes and greenstone belts. A key question relating

to the tectonics of greenstone belts is therefore the original

spatial relationship between the volcanic assemblages and

presumed-basement gneisses, and how this relationship has been

modified by subsequent deformation. Unconformities have been

reported from the Cameron River Greenstone Belt northeast of

Yellowknife, and from the Point Lake area to the north (1,2).

What remains unclear in these examples is the significance of

the so-called "later faulting" of the greenstone - gneiss

contacts. Do the angular discordances really represent

unconformities, or could they be better - interpreted as a

consequence of the juxtaposition of originally widely separated

terranes? Where unconformities between gneisses and overlying

sediments are indisputable, such as at Point Lake, the

significance of faults which occur below the base of the

volcanic succession also needs to be evaluated. As part of an

on-going investigation aimed at answering these and other

questions, I mapped the extremely well-exposed Cameron River

Greenstone Belt and the Sleepy Dragon Metamorphic Complex in the

vicinity of Webb Lake and Sleepy Dragon Lake during the summer

of 1985, extending the efforts of earlier workers (3,4,5,6).

The greenstone belt was found to consist predominantly of

mafic pillowed to massive flows and numerous dike complexes. At

the preserved base of the greenstone belt these dikes locally

retain a sheeted aspect and display one-way chilling.

Subordinate amounts of pyroclastic rocks and volcanic breccias

are also present. Rocks of the Sleepy Dragon Metamorphic

Complex are highly variable, and include both ortho- and

paragneisses, along with numerous mylonite zones (3,7,8). Older

gneisses and mylonites are intruded by several younger phases of

mafic to silicic plutonic rocks which show different intensities

ofj^def ormation.

The contact between the Cameron River Greenstone Belt and

the Sleepy Dragon Metamorphic Complex was found to be a

half-kilometer wide zone of very complex structure. All rocks

within this high-strain zone have a strong steeply plunging

stretching lineation, although rocks from throughout the area

also have a less-intense generally vertical lineation.

Transposed layering and intensely folded quartz segregations are

common in this zone; sheath folds with verticaly plunging hinges

are present in some localities, indicating very high shear

strains. Macroscopic sense-of-shear indicators are not abundant

but generally suggest that the Cameron River Belt was thrust

over the Sleepy Dragon Complex. Supporting microscopic work is
70 -

currently underway. In one area north of Webb Lake some slivers

of gneissic basement are intercalated with phyllonites of the

major high-strain zone; further mapping will reveal the lateral

extent of this tectonic juxtaposition, but it is apparently the

first-documented example of Archean basement-involved thrusting

in the Slave Province.

Pillow lavas of the Cameron River Belt immediately adjacent

to the basal high-strain zone have aspect ratios locally

exceeding 1:3:10. In an area to the northwest of Sleepy Dragon

Lake these lavas are overturned in an anticline as shown by

locally-consistent facing directions. The axial trace of this

fold is parallel to the contact zone, and the fold's geometry is

consistent with formation during thrusting of the Cameron River

Belt over the Sleepy Dragon Complex. Preliminary mapping in the

greenstone belt in the Webb Lake area has revealed the presence

of a few other subparallel shear zones containing structures

similar to those just described; a common origin is tentatively

inferred pending more detailed mapping.

Interpreting the structures within the Sleepy Dragon

Metamorphic Complex is difficult because of the complex

deformation history of this terrane. The only structure which,

at this point, can unambiguously be related to movement along

the contact with the Cameron River Belt is a foliation which

trends parallel to and increases in intensity towards the

contact zone. The foliation cuts earlier structures including

folded gneissic and mylonitic foliations; earlier foliations are

folded about this later one (3). The fact that this late

foliation is cut by some plutonic bodies suggests that a minimum

age may be placed on the thrusting and emplacement of the

Cameron River Greenstone Belt over the Sleepy Dragon Metamorphic

Complex.

Numerous mafic dikes are present both at the base of the

Cameron River Belt and within the Sleepy Dragon Complex near

it's contact v/ith the greenstone belt (6). The textures and

xenolith content of the dikes in the Sleepy Dragon Complex

appear to be generally different from the dikes in the

greenstone belt. Deformational and metamorphic fabrics in the

dikes of the Sleepy Dragon Metamorphic Complex suggest that they

are of at least two, and probably three generations, while only

two distinct generations of dikes are recognized from the

Cameron River Greenstone Belt. Pending further field and

laboratory work it is tentatively suggested that (a) the first

two generations of dikes in the Sleepy Dragon Complex are not

directly related to any dikes in the greenstone belt, (b) the

earliest generation of (locally sheeted) dikes in the greenstone

belt are not present in the basement complex, and (c) only the

latest, relatively undeformed dikes are correlatable between the

two terranes.

Although it is a bit premature to propose tectonic models

for the Cameron River Greenstone Belt it is useful to keep a

working hypothesis in hand. It is tentatively proposed that the

Sleepy Dragon Complex is a preserved remnant of a rifted Archean

continent; many of the metasedimentary gneisses may represent a
71

highly deformed Atlantic-type margin sequence originally

deposited on top of older basement. The Cameron River

Greenstone belt has many affinities with Phanerozoic ophiolites

and/or island arc complexes, including the presence of sheeted

dikes. Although an unconformable relationship has been reported

between the Cameron River Belt and the Sleepy Dragon Complex, I

have not yet been able to support this contention based on

observed field relationships. In fact, all data collected

to-date indicates that the greenstone belt is allochthonous.

Structures at and near the base of the greenstone belt suggest

that it has been imbricated and thrust over the Sleepy Dragon

Metamorphic Complex, incorporating slices of gneiss in the

process. It must be emphasized that these are only preliminary

conclusions that need to be verified by several more seasons of

detailed mapping but, so far, many similarities are seen between

Slave Province greenstone belts and Phanerozoic collisional

tectonic zones.

ACKNOWLGEMENTS

Funding was provided by the Geological Society of America

and by The Johns Hopkins University.

BIBLIOGRAPHY

(1) Baragar, W.

(2) Henderson,

Plate Tectonics

(3) Davidson,

G.S.C., Pap. 72

(4) Lambert,

October, 1973,

(5) Lambert, M.

340

(6) Lambert, M.

(7) Henderson,

(8) Kusky, T.

press

R., and J.C. McGlynn, 1976, G.S.C., Paper 76-14

J.B., 1981, in: A. Kroner (ed.), Precambrian

, pp. 213 - 235

A., 1972, in: Report of Activities, Part A,

-la, pp. 109 - 115

M.B., 1974, in: Report of Activities, April to

G.S.C., Pap. 74-1, Pt. A, pp. 177-179

B., 1977, G.A.C., Special Paper no. 16, pp. 331 -

B., 1982, G.S.C., Paper 82-1A, pp. 165 - 167

J.B., 1976, G.S.C. Open File 353

M., 1986, Cont. Geol. of the N.W.T., vol. 3, in
72 N86-23114

SEDIMENTOLOGICAL AND STRATIGRAPHIC EVOLUTION OF THE SOUTHERN

PART OF THE BARBERTON GREENSTONE BELT: A CASE OF CHANGING PROVENANCE AND

STABILITY; Donald R. Lowe and Gary R. Byerly, Department of Geology,

Louisiana State University, Baton Rouge, Louisiana 70803 USA

The sedimentological and stratigraphic evolution of the 3.5 to 3.3 Ga

Barberton Greenstone Belt can be divided into three principal stages: (1)

the volcanic platform stage during which at least 8 km of mafic and

ultramafic volcanic rocks, minor felsic volcanic units, and thin sedimentary layers (Onverwacht Group) accumulated under generally anorogenic

conditions, (2) a transitional stage of developing instability during

which widespread dacitic volcanism and associated pyroclastic and volcaniclastic sedimentation was punctuated by the deposition of terrigenous

debris derived by uplift and shallow erosion of the belt itself (Fig Tree

Group) , (3) an orogenic stage involving cessation of active volcanism,

extensive thrust faulting, and widespread deposition of clastic sediments

representing deep erosion of the greenstone belt sequence as well as

sources outside of the belt (Hoodies Group).

I. The platform stage of Barberton Greenstone Belt development is

represented by rocks of the predominantly volcanic Onverwacht Group.

Sediments deposited during this stage included (a) dacitic breccias,

conglomerate, and coarse sands deposited as part of and adjacent to felsic

volcanic centers and, less abundantly, proximal mafic lapillistones and

tuffs; (b) distal felsic volcaniclastic and pyroclastic layers consisting

.mainly of fine ash, dust, and accretionary lapilli, (c) biogenic deposits

such as carbonaceous oozes, carbonaceous muds, bacterial mats, and locally,

stromatolites, and (d) orthochemical sediments including evaporites,

barite, carbonate, and possibly siliceous deposits. The bulk of these

sedimentary units show clear evidence of having been deposited under

shallow-water conditions. The regional stratigraphic continuity and

sedimentological integrity of sedimentary layers within this sequence, the

predominantly shallow-water depositional setting, and the paucity of

debris derived from the uplift and erosion of older rock sequences indicate

that the overall depositional and tectonic setting was a broad, low-relief,

shallow-water anorogenic platform (1).

II. Rocks traditionally assigned to the Fig Tree Group were deposited

during a transitional phase of greenstone belt evolution. These are

exposed in a complex succession of thrust sheets that provide numerous

exposures of each part of the stratigraphic sequence (2). The lowest part

of the Fig Tree is characterized by distal volcaniclastic units and

carbonaceous cherts resembling those in the Onverwacht but showing rapid

lateral facies changes. In particular, 40 to 50 m of predominantly

carbonaceous chert in some structural belts can be correlated with a

sequence of interbedded ultramafic lavas, banded cherts, carbonaceous

cherts, stromatolites, and volcaniclastic units at least 500 m thick in

other areas (2).

The overlying 200 to 500 m of rocks includes two principal components.

By far the greatest thicknesses of Fig Tree strata consist of heavily

altered dacitic pyroclastic and volcaniclastic detritus (3). This succession includes three main lithofacies: (a) plagioclase-phyric intrusive

rocks that may locally grade into extrusive flows, (b) proximal, plagioclase-phyric breccias and conglomerates, probably developed as lava domes

and surrounding coarse epiclastic units, and (c) regionally extensive ash
BARBERTON GREENSTONE BELT

Lowe, D. R., and Byerly, G. R. 73

deposits, tuffs, and their current-worked equivalents, volcaniclastic

sandstone and siltstone. The bulk of the finely laminated cherty ferriginous sediments characterizing Fig Tree rocks throughout much of the

Mountain Land represent altered fine-grained dacitic volcaniclastic

deposits. In contrast to previous interpretations, we consider the Fig

Tree to represent a predominantly volcanic interval, perhaps more closely

related petrogenetically to the Onverwacht Group than to the suprajacent

erogenic Hoodies succession.

Interbedded with these volcanic and volcaniclastic strata are thin,

lenticular units of chert-pebble conglomerate and chert-grit sandstone

showing rapid lateral facies changes and apparently representing debris

derived from local uplifts within the greenstone belt. Most of the debris

can be identified with underlying silicified rocks of the Fig Tree Group;

there is little evidence for major uplift or deep erosion of the greenstone belt at this time.

III. Rocks which have traditionally been included within the Moodies

Group represent three main clastic lithofacies: (a) a sequence of quartzpoor, highly altered sands and fine gravels derived by erosion of the

subjacent dacitic rocks; (b) thick, coarse, chert-clast conglomerate and

chert-grit sandstone derived by weathering and erosion of uplifted parts

of the greenstone belt, and (c) quartzose and locally K-spar-rich sandstone

representing the erosion of sources outside of the greenstone belt,

possibly but not necessarily including the intrusive granitoid rocks

and/or the Ancient Gneiss Complex or its equivalents.

Although the stratigraphic sections in most structural belts can be

correlated with one another, there is as yet no satisfactory reconstruction

of their original relative depositional positions. So-called northern

facies rocks in the Mountain Land also belong to allochthonous terranes

and their present location relative to units to the south is clearly of

tectonic rather than depositional origin.

The overall sequence includes numerous minor unconformities and at

least one major break. Within the Onverwacht Group, pauses in effusive

activity are marked locally by weathering and erosion of flow surfaces,

but no significant formation or accumulation of clastic debris. The

inception of felsic volcanism both in the upper Hooggenoeg formation and

the Fig Tree Group was accompanied by minor instability and local erosion

of underlying rocks. Also, the formation of large, high-relief subaerial

felsic volcanic edifices in Hooggenoeg and Fig Tree times was followed by

extensive erosion and truncation of these complexes. The major structural

unconformity within the Barberton sequence occurs locally at the base of

the Moodies Group. Although a number of apparently conformable Fig

Tree-Moodies transitions occur, over wide areas, the Moodies was deposited

with angular unconformity on rocks as old as the Hooggenoeg Formation.

This contact has additionally been complicated by structural movement.

The sedimentological development of the Barberton Greenstone Belt

reflects three principal tectonic stages involving three contrasting

sources of clastic sediment. The volcanic platform stage, represented by

rocks of the Onverwacht and Fig Tree Groups, was primarily an interval of

rapid effusion of lavas, subsidence, but little differential tectonic

movement. The main sources of clastic detritus were first cycle, active,

high-relief, felsic and, to a lesser extent, mafic volcanic centers. The

second stage, represented by rocks of the Fig Tree Group, was one characterized by continuing, regionally extensive volcanism and developing
BARBERTON GREENSTONE BELT

74 Lowe, D. R., and Byerly, G. R.

tectonic Instability reflected by the presence of extensive lateral facies

changes and small intra-platform uplifts that supplied shallow-level

intrafotmational debris to local sedimentary systems. Latest Fig Tree and

Hoodies deposition was influenced by concurrent thrusting and orogenesis.

Sediments were derived initially from both shallow and deep levels within

the greenstone belt and, later, from distant quartz and K-spar rich

sources outside of the belt.

REFERENCES: 1. Lowe, D. R. (1982) Precam. Res., 17, 1-29. 2. Lowe, D.

R., et al. (1985) Precam. Res., 27, 165-186. 3. Byerly, G. R., and Lowe,

D. R. (1984) Proc. Lunar Planet. Sci. Conf. 16th, p. 101-102.
ORIGINAL PAGE IS 75

OF POOR QUALITY fl 86 -23 US

EVIDENCE FOR STRUCTURAL STACKING AMD REPETITION IN THE

GREENSTONES OF THE KALCOORLIE DISTRICT. WESU.RN AUSTRALIA

J.I. Nartyn, Esso Australia l.hi., Sydney

INTRODUCTION

Host previous stratlgraphlc Interpretations of the southern part of the Norseman -Wi luna

Greenstone Belt have proposed polycycllc sequences (e.g. Itorwitt and Sofoulls'; McCall^;

Williams'; Gllkson4; Gamuts and Theron^). These Invoked two and sometimes three

successive suites of mafic and/or ultramaflc volcanlcs and intrusives separated by felsic

volcanics and immature clastic sediments, however no distinctive lithological differences were

reported between successive maf ic -ultramaf Ic sequences. When interpretations of Williams et

a 1.6 and Hal Iborg^, further to the north, are integrated, a total of four separate major

mafic-ultramafIc suites emerges for a large part of the Norseman-Wiluna Belt. Although the

author does not intend to imply that all polycycllc stratigraphies are wrong in principle luch

a situation seems suspiciously over-complex and stimulates the need to look critically at the

individual areas where stratigraphies have been erected. For the Kalgoorlie area in the south,

some of the schemes have already provoked scepticism (Burke et al.&; Archibald et at.*) and

a simpler model consisting of one cycle subject to structural repetition has been evolved by

workers in the Geological Survey of Western Australia (Griffin et al.'°) for part of this

area. The latter authors drew attention to the 'carbon copy' similarity between the elements of

some polycycllc stratigraphies. Much more regionally extensive integrated structural and

stratlgraphic data is still required to evaluate the relationship between structure and

stratigraphy more fully, an objective substantially limited by poor outcrop and deep

weathering, but with due effort, far from unattainable.

OUTLINE OF STUDY

Regional mapping by the author in an area of approximately 20,000 km2 centred on Kalgoorlie

revealed many problems and anomalies In several of the published stratigraphic schemes.

However since insufficient critical stratlgraphlc and structural evidence had been given in

support of the schemes it has not been easy to check the bases on which they were erected. The

following lines of investigation have been pursued.

* Regional distribution and interelationships of lithologically similar sequences previously

regarded as distinct, based on mapping, mineral exploration data, and geophysical

interpretation. Emphasis has been on the mafic-ultramafic suites because they are the most

easy to define and map.

* Critical evaluation of contacts and their associated structural features.

RESULTS

There are several instances where mafic-ultramaf ic suites previously proposed as younger (e.g.

Coolgardie-Kurrawang area in Glikson*) join or merge with their 'older* counterparts when

mapped over various distances. They range in size from splinter-like splays a few kilometres

long diverging from a major mafic belt by up ho a kilometre, to extensive sheets which are

traceable for tens of kilometres as separate entities before joining with and becoming

indistinguishable from their 'older' counterparts. Some successions are isolated in

metasedimentary terrain, and never connect with their sequences of origin; however this

situation is unusual. In areas where like elements of two proposed cycles are juxtaposed or

interconnected (e.g. Widgiemooltha and Spargovilie areas in map of Gemuts and Theroo5) there

seems to be no clear reason why they should have been regarded as separate.
„* t fJ *' 'EVIDENCE;,FWI STRUCTURAL REPETITION

76 " r/J.E. 'MARTYlP

The apparent stratigraphic thicknesses of many of the previously proposed younger

mafic-ultramafic sequences is very variable. While they nay be measurered in kilometres in some

areas, in many localities the sequences are attenuated and deformed. They may be traced for

tens of kilometres as apparently conformable packages of all or most of the major mafic and

ultraoafic lithologies, though individually these lithologies may occur as lenses or sheets

hundreds or even only tens of metres in thickness. While such sequences have been interpreted

in the past as volcanic intercalations in a eugeosynclinal sedimentary pile (McCall2;

Glikson4), or the beginnings of new volcanic cycles, their degree of deformation and tendency

to be smaller scale carbon copies of their 'older* counterparts is more consistent with

structural repetition. In a number of instances they are overlain by felsic volcanic rocks

suggesting cyclic development in a uniformly facing sequence. This is here regarded as evidence

that repetition has been mainly by faulting and not by recumbent or isoclinal folding.

OBSERVATIONS OH CONTACTS

Many previous stratigraphies (e.g. Williams'; Gemuts and Theron^; Glikson4) have been

erected in areas where fragmentary facing evidence suggests thick uniformly facing sequences.

The potential for strike dislocations has generally been overlooked despite heterogeneous shear

deformation. There has been an absence of critical treatment of major formations I contacts to

establish whether they are normal or tectenised. This is understandable in some instances

since such contacts are rarely well exposed, however diligent search by the author has revealed

many key outcrops. The vast majority of these provide compelling evidence that all is not well

with the published polycyclic stratigraphies. Examination of contacts, especially basal ones,

and the contact areas of the previously proposed younger mafic-ultramafic suites, commonly

reveals strong peneconcordant shearing, recrystellised mylonitic or other cataclastic rocks, or

in one instance (the Kalpini formation which is the highest mafic-ultramafic suite of

Williams'), an overturned but undeformed contact with clear facing evidence the reverse of

that previously proposed. Napping the relationship between the contact zones and primary

layering often reveals subtle discordances not readily explained by unconformity.

MODEL

Many proposed structural repetitions or fault slices are linear, others are arcuate and folded

around major upright structures. Linear belts are often controlled by throughgoing

transcurrent deformation zones with pronounced sub-horizontal Iineations. Arcuate systems

however were conceivably generated by earlier processes such as thrusting or gravitational

gliding predating upright folding. Although transcurrent shearing is a feasible mechanism for

repetition for at least some of the more linear belts, it is possible that even many of these

began life as early thrust sheets and became stretched and aligned by later transcurrent

deformation. Early thrusts, recumbent folds and layer-parallel shear fabrics have been

documented in several localities in the Norseman-Wiluna Belt where prevailing strikes deviate

from the NNW regional grain, or where tight upright folding is subdued or absent (e.g.

Chapman11; Gresham and Loftus-HiI Is12; Archibald1'; Platt et at.14; Nartyn and

Johnson'^; Spray16). In one instance (Chapman") a narrow mafic-ultramafic belt in

sediments has clearly been generated by an overthrust. Almost certainly the recognition of

this structural style in east-west trending or gently domed areas is a consequence of

preservation. It is undoubtedly present also in NNW trending linear domains but is overprinted

and hard te recognise. It is emphasised that thrust repetition does not explain all of the

previously proposed younger cycles in the district. Some are a consequence of

misinterpretation, by placing too great a significance on isolated stratigraphic facing

observations, or from attempts te correlate across major upright faults. Broad regional

observations by the author suggest that thrust repetition may be much more strongly developed

in the Kalgoorlie district than elsewhere in the Norseman-Wiluna Belt though this conclusion is

tentative.
• *j. .* ORIGINAL PAGE BS

EViOENCB%OR STRUCTURAL REPETITION OF POOR QUALITY

J.E. MARTYN 77

Thrusting does not- appear to have occurred on a scale comparable with many Phanerozoic

convergent plate boundaries. There Is no evidence of juxtaposition of strongly contrasting

domains, or of high pressure metamorphism. There is also a lack of pronounced east-west

assymetry across the Norseman-Wiluna Belt as a whole. The tectonics can be viewed more in term

of a rearrangement of familiar elements of the local stratigraphy, a situation more consistent

with a closed or intracratonic setting, rattier than an open plate margin. This accords with

models such as those of Groves'7. As such, intrabasinal gravity gliding resulting from early

uplift heralding later vertical tectonic events is the most favoured model by the author. This

is consistent with the sedimentation style which is dominated by turbidites and includes debri

flow deposits. 01Istostromes have also been reported (Taylor, in Gee and Groves '"). In some

respects the scheme resembles that proposed by De Wit'9 for the Barberton Greenstone Belt.

Felsic volcanism was intimately associated with sedimentation, and It Is possible that

concomitant granitic intrusion Into a dense sheet of mafic-ultramafic volcanics may have

triggered the instability that first led to the sedimentation and later to gravity gliding

tectonics. Subsequent folding and faulting of the tec too leal Iy stacked sequence would have

created the illusion of a polycyclic sequence which has suffered only upright folding and

shearing. The upright tectonic events have generated their own set of interpretive problems.

Peneplenation, and Tertiary lateritic weathering ultimately obscured much of the important

evidence.

REFERENCES

1. Horwitz, R.C., andSofoulis, J. 1965. Proc. Aust. Inst. Mining Metal I. 214, p. 43-59.

2. McCall, G.J.H., 1969. Proc. Roy. Soc. West. Aust., 52, p.119-128.

3. Williams, I.R., 1970. Geol. Surv. West. Aust. Rec. No. 1970/1.

4. Glikson, A.Y., 1971. Spec. Publ. No.3, Geol. Soc. Aust. p.443-460.

5. Gamuts, I. and Theron, A.C., 1975. Aust. Inst. Mining Metal I. Hon. 5, p.66-74

6. Williams, I.R., Gower, C.F., and Thorn, R., 1971. Gaol. Surv. West. Aust. Rec.No.1971/26.

7. Ha11berg, J.A., 1983. Gaol. Surv. West.Aust. Rec. No. 1983/8.

8. Burke, K., Oavey, J.F., and Kidd, W.S.F., 1976. The early history of the Earth. John Wiley

and Sons, p. 113-129.

9. Archibald, N.J., Bettany, L.F., Bickle, M.J. and Groves, D.I., 1981 Spec. Publ. No. 7 Geol.

Soc. Aust. p.491-504.

10. Griffin, T.J., Hunter, W.M., and Keats, W. 1983. Eastern Goldfields Geological Field

Conference, Excursion and Guide. Geol. Soc. Aust. p.7-8.

11. Chapman, O.M., 1982. B.Sc. Hons. Thesis, Univ. of West. Aust.

12. Gresham, J.J. and Loftus-HiI Is, G.D., 1981. Econ. Gaol. 76, p.1373-1416.

13. Archibald, N.J., 1979. Ph.D. Thesis, Univ. of West. Aust.

14. Platt, J.P., Allchurch, P.O. and Rutland, R.W.R., 1978 Precamb. Res. 7, p.3-30

15. Martyn, J.E. and Johnson, G.I., in press. Aust. J. Earth Sciences.

16. Spray, J.G., 1985. J. Struct. Geol. 7(2) p.187-203.

17. Groves, D.I., 1982. Revista Brasiliera da Geosciencias, 12, p.135-148.

18. Gee, R.O., and Groves, O.I. 1982. Univ. of West. Aust. Geol. Oept. and Extension Service

Publ. No. 7, p.C3-€IO.

19. Oe Wit, N.J., 1982. J. Struct. Geol. 4, 117-136.
78

N86.23116

PRELIMINARY STRUCTURAL MODEL FOR THE SOUTHWESTERN PART OF

THE MICHIPICOTEN GREENSTONE BELT, ONTARIO; George E. McGill and

Catherine H. Shrady, Dept. of Geology and Geography, University of

Massachusetts, Amberst, HA 01003

The southwestern part of the Michipicoten Greenstone Belt includes

a 100 km2 fume kill extending northeastwards from the town of Wawa,

Ontario. Except for a strip along the Magpie River that is covered by

Pleistocene gravels, outcrop in the fume kill averages about 30-507o.

Within this area are all the major lithologic belts characteristic of

the southwestern fourth of the Michipicoten Greenstone Belt. All of

the area mapped to date lies within Chabanel Township, recently mapped

at 4" = 1 mile by Sage et al. (1). Following a brief reconnaissance

in 1983, mapping at a scale of 1" = 400' was begun within and adjacent

to the fume kill in 1984. We have concentrated on two objectives:

1) determination of the geometry and sequence of folding, faulting,

cleavage development, and intrusion; and 2) defining and tracing lithologic "packages", and evaluating the nature of the contacts between

these packages. Results for objective 1) are discussed in a companion

abstract (2); this abstract will present tentative results for objective

2).

The entire Michipicoten Greenstone Belt has experienced relatively

late movement on steep faults, most of which trend approximately NNW

or NE (1,2,3). Some of this movement preceded the emplacement of diabase

dikes, some followed. These displacements may be easily removed in

order to reassemble older structures, which are of much greater tectonic

interest.

For mapping and descriptive purposes, it long has been customary

to divide the stratified rocks of the Michipicoten Greenstone Belt

into 4 major lithologic groups (1,3): mafic-intermediate volcanics,

intermediate-felsic volcanics, clastic sediments, and chemical sediments

(including iron formation). This is certainly valid, because outcrop

belts of these groups maintain integrity for long distances. However,

there are along-strike intergradations among them, and there is no

easy way to correlate between physically separated belts of similar

lithology. This last problem means that there is no really dependable

belt-wide stratigraphy, and relative ages of the various belts of similar

lithology are known only in the few places where modern radiometric

ages have been measured (4,5).

Our detailed mapping (Fig. 1) indicates that the situation is

more complex than one would infer from published maps and descriptions

(1,3,6). There are several lithologic packages within the single belt

of clastic sediments in Chabanel Township, all of which appear to be

bounded by fault contacts. In some cases, stratigraphic way up reverses

across these faults, in other cases it does not. At map scale, the

package boundaries follow bedding or volcanic layering on one or both

sides, but locally this is not so, and at outcrop scale it commonly
STRUCTURAL MODEL, MICHIPICOTEN BELT

McGill, G.E. and Shrady, C.H. 79

ORIGINAL PAGE IS

OF POOR QUALITY

|T\ ?"' ^\ \ N

— t«on tormoton

...- Contoct bctwMA BvprocrutloU

Ond granitic rock*

I 1 GroniK l«ch«

MtCHIPICOTEN GREENSTONE BCLT. ONTARIO

E«plo notion

P o ?L^ Pleistocene grovels

C-X'X'XI Coorse. polymiclic voteomc-closl conglomerolt

U-I-I-5-1 Oommontly orgillile

Croywocke ond orgillile (turbidile)

Faults bounding hthologic packages

Other faults

— • -Temporary limit of mopping

"SftFwvsSi rv

^.••:/:-'-.:'£>!4 /-Q ^ — '• ;

T / L"

)0

t ^ Lo(l«

't-L^ i~_,

^

Graywacke, orkos* and quortiile

^*>I*j»l Pyroclostici ond volcaniclastic*

T

,',',','\ Inlermediau-felsic volcanic*

*.*.*.*J Mafic-intermediate volconics

• Contacts

-*- -rf- Strike ond dip; upright, overturned

Fig. 1. Geologic sketch map of the central part of Chabanel Township,

Ontario. All intrusive igneous rocks omitted for simplicity.

B-B1 and A-A1 indicate corresponding points across late faults.
STRUCTURAL MODEL, MICHIPICOTEN BELT

80 McGill, G.E. and Shrady, C.H.

is not so. In places, these faulted boundaries are characterized by

locally developed cleavages, excessive flattening or elongation of

pebbles, or minor folds.

The area we have mapped seems to be a zone of faults and folds

separating a large region to the south underlain by overturned rocks

with tops north from an even larger region to the north underlain by

overturned rocks with tops south (1,6). This relationship would seem

to indicate an antiformal fold in the inverted limb of a very large

nappe, but we have not been able to define such a structure, and rocks

that should correlate across the structure are not the same age (R.

Sage, pers. com.). Major faulting thus is necessary, but earlier or

synchronous folding at township or larger scale would seem necessary

to account for the opposed overturning. Almost all of the rocks north

and south of our area are volcanic, so it may never be possible to

determine if these terranes consist of continuous sections or if they,

too, are divided into fault-bounded packages.

Because we have yet to sort out the sequence of minor and major

structures with sufficient confidence, and because completed detailed

mapping covers such a small fraction of the total belt, we prefer to

be rather conservative about interpreting our data. Key observations

include a "stratigraphy" that consists mostly of fault-bounded

"packages", the apparent early age of these faults, and the large areal

extent of the inverted sequences facing each other. The most attractive

and probably the simplest explanation for these relationships involves

early imbricate thrusting—before the imposition of the almost universal

steep dips. However, this interpretation remains to be proved.

References

(1) Sage, R.P., et al. , (1982) Ontario Geol. Surv. Prelim. Maps of

Precambrian Geology: a. Chabanel Twp., Map P. 2439; b. Esquega

Twp., Map P. 2440; c. McMurray Twp., Map P. 2441; d. Lastheels

Twp., Map P. 2442; e. Musquash Twp., Map. 2568.

(2) Shrady, C.H. and McGill, G.E. (1986) This volume.

(3) Sage, R.P., et al. , (1984) Prelim. Map of Precambrian Geology

of Corbiere Twp., Ontario Geol. Surv. Map P. 2607.

(4) Turek, A., et al., (1982) Can. Jour. Earth Sci., 19, 1608-1626.

(5) Turek, A., et al., Can. Jour. Earth Sci., 21, 457-464.

(6) Goodwin, A.M. (1962) Geol. Soc. America Bull., 73, 561-586.
N86-23117 81

THERMAL IMPLICATIONS OF METAMORPHISM IN GREENSTONE BELTS AND THE HOT

ASTHENOSPHERE-THICK CONTINENTAL LITHOSPHERE PARADOX; Paul Morgan, Department

Geosciences, Purdue University, West Lafayette, IN 47907.

From considerations of secular cooling of the Earth and the slow decay of

radiogenic heat sources in the Earth with time, the conclusion that global

heat loss must have been higher in the Archean than at present seems

inescapable. The mechanism by which this additional heat was lost and the

implications of higher heat loss for crustal temperatures are fundamental

unknowns in our current understanding of Archean tectonics and geological

processes. Higher heat loss implies that the average global geothermal

gradient was higher in the Archean than at present, and the restriction of

ultramafic komatiites to the Archean and other considerations suggests that

the average temperature of the mantle was several hundred degrees hotter

during the Archean than today (1). In contrast, there is little petrologic

evidence that the conditions of metamorphism or crustal thickness (including

maximum crustal thickness under mountains) were different in Archean

continental crust from the Phanerozoic record (see 1). Additionally, Archean

ages have recently been determined for inclusions in diamonds from Cretaceous

kimberlites in South Africa (2), indicating temperatures of 900 to 1300 degC

at depths of 150 to 215 km (45 to 65 kbar) in the Archean mantle (3), again

implying relatively low geothermal gradients at least locally in the Archean.

In this contribution the thermal implications of metamorphism are examined,

with special reference to greenstone belts, and a new thermal model of the

continental lithosphere is suggested which is consistent with thick

continental lithosphere and high asthenosphere temperatures in the Archean.

High-grade metamorphism is common in Archean terrains (4, 5), and

includes some greenstone belts, such as in the Yilgarn block of SW Australia

(6). High metamorphic temperatures (700 degC or more) and often high

metamorphic pressures (5 to 10 kbar or greater) are indicated by the mineral

assemblages in these terranes, and they are underlain in most cases by

continental crust of normal thickness (7, 8). Conductive thermal relaxation

models have been proposed to predict the thermal conditions of metamorphism in

the crust following tectonic activity such as underthrusting (e.g., 9-11). As

demonstrated by Ashwal and Morgan (7), however, simple thermal relaxation of

thickened crust cannot reasonably produce the high temperatures required by

granulite metamorphism with a thick section of crust (30 km or more) below the

shallowest depth of granulite metamorphism without requiring the lower part of

the crust to be supersolidus. Basically the temperature range for granulite

metamorphism is so close to estimates of the crustal solidus for reasonable

crustal compositions (e.g., 12), that a positive geothermal gradient below the

shallowest depth of granulite metamorphism causes the geotherm to intersect

the solidus above the Moho. Ashwal and Morgan (7) conclude that unless

granulite metamorphism occurs only near the base of the crust and the thick

section of crust now below the exposed granulites was added after

metamorphism, major crustal magmatic activity is associated with granulite

metamorphism. Such extreme thermal conditions are not required by lower

grades of metamorphism, but any metamorphic gradients which indicate a high

geotherm suggest the upward transport of heat by magma unless the crust is

thin.

If it is accepted that magmatic heat transport is an essential component

of the crustal thermal regime during the peak thermal conditions recorded by

the metamorphic mineral assemblages in the crust (at least where high

geothermal gradients are indicated), then maximum temperatures recorded in
- f ' Tt.«

o -A „„. O ?1

GREENSTONE BELTS AND GEOTHERMS

82 Morgan, P.

these systems were buffered by the solidus. The occurrence of young

granulites at the top of sections of normal thickness crustal sections

similarly indicates that modern maximum geothermal gradients are buffered by

the solidus. A similar conclusion is indicated by heat flow data from areas

of recent tectonism in which high heat flow must result from magmatic heating

of the crust (e.g., 13). Maximum temperatures at shallow depth are buffered

by the boiling point curve at hydrostatic or lithostatic pressures, below

which maximum temperatures are buffered by the crustal solidus. As these

maximum crustal temperatures are commonly encountered in areas of active

tectonism and magmatism today, it is impossible for maximum temperatures

recorded by Archean metamorphic assemblages to have been higher than modern

maximum temperature conditions unless the solidus was different. Thus, in

this buffered system, higher heat loss in the Archean is not expected to be

recorded by metamorphic assemblages indicating higher geothermal gradients

than peak modern conditions, although these peak crustal thermal conditions

may have been more widespread in the Arcbean than at present.

The occurrence of high-grade (granulite) metamorphism in Archean

greenstone belts suggests that either the high-grade areas were produced near

the base of the crust and subsequently the crust has been thickened below the

high-grade terranes, and/or magmatism was an important process during the

high-grade metamorphism. The intimate association of plutons with the

greenstone belts in "granite-greenstone" terranes suggests the importance of

magmatism during this high grade metamorphism, and is consistent with models

which suggest basal melting of stacked simatic thrust sheets during the

evolution of at least some greenstone belts (14-16).

Perhaps the most paradoxical indicator of Archean thermal conditions with

respect to higher global heat loss is the relatively low Archean geothermal

gradients indicated by the formation of diamonds of Archean age. The diamond

stability field is consistent with geotherms predicted for modern shield areas

with thick (150 km or greater) lithosphere (e.g., 13). Meyer (3) has

suggested that diamonds were formed in the asthenosphere which in turn

suggests that perhaps the higher temperatures deduced for the Arcbean mantle

from the occurrence of komatiitic lavas were not universal. A more common

interpretation of the diamond data is that they indicate the existence of

thick "keels" of subcontinental lithosphere below at least some areas during

the Archean (1, 16). However, as the lithosphere is intimately related to the

thermal boundary of upper mantle convection, it would be expected that this

boundary layer and the lithosphere would have been thinner during the Archean

with higher global heat loss and mantle temperatures. A possible solution to

this paradox may be found in the intrinsic heat production of continental

lithosphere.

There are two basic variable parameters that control the stable thickness

of the continental thermal boundary layer (lithosphere), the heat production

within the layer and the heat input to its base (13, 17). The layer thins if

heat input to its base increases, and thickens if the heat input decreases.

This heat input depends upon the temperature difference between the lower

portion of the stable boundary layer and the underlying convection cell, or

more specifically the temperature gradient in the lowest portion of the layer.

As this gradient decreases to zero, the heat input to the base of the

lithosphere decreases to zero (negative gradients are not permissible in a

stable thermal boundary layer). The thickness of stable continental

lithosphere with zero heat input at its base is independent of the global heat

loss, assuming that the heat can be lost elsewhere (oceanic and other

continental lithosphere), and this may possibly be a mechanism for maintaining
I "GREENSTONE BELTS AND GEOTHERMS

Morgan, P. 83

thick continental lithospbere at a time of high global heat loss and high

average mantle temperatures.

The condition for zero heat flux into the base of the stable continental

litbosphere is that the temperature increase within the lithosphere due to its

intrinsic radiogenic heat production creates a geotherm that is asymptotic to

the asthenosphere isotherm (or adiabat with an adiabatic basal heat flux).

For thick lithosphere this condition requires a small but significant

component of heat production in the mantle lithosphere, and an example of such

a heat production distribution and geotherm are given in Figure 1. This

condition has the interesting property that thicker lithosphere is indicated

for higher asthenosphere temperatures for similar heat production

distributions. If heat production distributions of this type are realistic it

is unlikely that they are accidental (see also 18), and the concentration of

radiogenic heat production into the lithosphere by metasomatism and crustal

building processes may be related to the stabilization of continental

lithosphere.

Figure 1. Example of continental

lithosphere geotherm asymptotic

with asthenosphere isotherm as a

result of its intrinsic radiogenic

heat generation. A two component

crustal heat generation model is

assumed for this geotherm: An upper

crustal component decreasing

exponentially with depth from 2.7

^W/m*»3 at the surface with a

depth scale length of 7 km, and an

additional uniform component of

0.09 /iW/m"3 (geotherm model

modified from 19).

i

100

1

°200

300

TEMPERATURE (degC)

500 1000 1500

i i

Asthenosphere

Isotherm

References; (1) Richter, P.M. (1985) Earth Planet Sci, Lett. 73, 350-360. (2)

Richardson, S.H., Gurney, J.J., Erlank, A.J. and Harris, J.W. (1985) Nature

310, 198. (3) Meyer, H.O.A. (1985) Am. Mineral. 70, 344-355. (4) Condie, K.C.

(1981) Archean Greenstone Belts, Elsevier, Amsterdam. (5) Windley, B.F. (1982)

The Evolving Continents, Wiley, Chichester. (6) Gale and Klein (1981) in (5).

(7) Ashwal, L.D. and Morgan, P. (1984) Heat sources for granulite

metamorphism, unpublished manuscript. (8) England, P. and Bickle, M. (1984) J.

Geol. 92, 144-146. (9) Bickle, M.J., Hawkesworth, C.J., England, P.C. and

Athey, D.R. (1975) Earth Planet. Sci. Lett. 26, 13-28. (10) England, P.C. and

Richardson, S.W. (1977) J. Geol. Soc. Lond. 134, 201-213. (11) Thompson, A.B.

(1981) Terra Cognita 1, 11-20. (12) Wyllie, P.J. (1977) Tectonphysics 43,

41-71. (13) Morgan, P. (1984) Phys. Chem. Earth 15, 107-193. (14) McGregor,

V.R. (1979) in Trondjhemites, Daoites and Related Rocks (F. Barker, ed.),

Elsevier, Amsterdam, 133-147. (15) Barker, F., Arth, J.G. and Hudson, T.

(1981) Phil. Trans. Roy. Soc. Lond. A301, 293-303. (16) de Wit, M. (1985) What

the oldest rocks say (abs.), Workshop on the Earth as a Planet, LPI and GSA,

Orlando, Florida, Oct. 27, 1985. (17) Morgan, P. and Sass, J.H. (1984) J.

Geodynamics.1, 143-166. (18) Morgan, P. (1985) J. Geophys. Res. 90, C561-C570.

(19) Davies, G.F. and Strebeck, J.W. (1982) Geophys. J. R. Astr. Soc. 69,

623-634.
84 N86-23118

GEOCHEMICA L CHARACTERS AN D TECTONI C EVOLUTIO N O F TH E

CHITRADURGA SCHIST BELT: AN ARCHAEAN SUTURE (.?) OF THE

DHARWAR CRATON, INDIA. S.M.NAQVI, NATIONAL GEOPHYSICAL

RESEARCH INSTITUTE, HYDERABAD, INDIA.

The Chltradurg a schist belt extendin g for abou t 450 km in a

NS directio n and 2-50 km across, is one of the most prominen t

Archaea n (2.6 b.y.) tectoni c feature s o f th e India n Precambrla n

terrain , comprisin g abou t 2 t o 1 0 km thic k sequenc e o f

vo 1 cano s ed i men t a ry rocks. The basa l unit of this belt is

composed of an orthoquartzite-carbonat e facies, unlike many other

contemporar y greenston e belts o f th e Gondwan a land whic h begi n

wit h a basa l mafic-u1tramafi c sequence . Eight y percen t of the

belt is mad e up of detrita l and chemogeni c sediments, their

sucession commencin g with a poorly preserved quart z pebble basal

conglomerat e and current bedded quartzite s which, in turn , rest

on tonalitic gneisses, the latter having been furthe r remobilized

alongwit h the schist belt. Deposition of current bedded matur e

arenites indicate the existence of platforma l conditions near the

shor e line. Polymicti c graywack e conglomerates, greywackes,

shales, phyllites, carbonates, BIFs (oxide , carbonat e an d

sulfide ) BMF's (Bande d Maganes e Formations) an d cherts thu s

constitut e th e mai n sedimentar y rock s o f th e belt. Th e

polymicti c conglomerate s contai n debris o f rock s o f olde r

greenston e sequences, a s wel l a s a n abundan t measur e o f folde d

quartzites, BIF's and gneissic fragments which represent earlier

orogenies.

Four differen t types of greywacke s are recognised in the,bel t

fro m N t o S . Mos t o f thes e hav e bee n derive d fro m th e

surroundin g tonaliti c gneisses whic h contained older greenstone

sequence s a s enclave s o f variou s dimensions. However , th e

younge r sequences in the nort h contain debris fro m the intrabasin

volcanis m also. The K-granite s and gneisse s ar e foun d to be

progressively abundan t in the source area of these graywacke s as

indicated by the graniti c componen t of the debris of the younge r

graywacke s sequences. Their REE pattern s ar e characterize d b y

both positive and negative Eu anomalies, the latte r especially in

t h e interbedde d shale s wit h grewackes. Geochemistr y o f th e

graywacke s and chemogenic sediments thus indicat e their deeper

oceani c environmen t o f formation . Althoug h stratigraphi c

relation betwee n the shallow wate r and deepe r wate r sediments is

uncertain , th e basa l orthoquartzites-carbonat e sequence s

indicatin g platforma l environmen t perhap s represen t a facie s

change due to shallow wate r conditions along the shore line, and

the greywack e suite those of deeper wate r away fro m it. Simila r

facie s chang e is observe d i n th e BIF's fro m shallowe r oxid e t o

deeper sulfid e facies.

The ultramafi c rocks, mostl y found in the lowe r sections of

t h e belt , sho w pillo w structure s an d spinife x textur e an d ar e

komatiiti c i n composition . Th e mafic , intermediat e an d aci d

volcanic s ar e foun d a s detache d outcrop s i n presumabl y highe r
\. : it , CHIT-RAgURGA BELT 85

r ' NAQVI' S". m.

stratigraphic sections and show tholeiitic and calc alkaline

affinities, probably produced by 5-15% melting. The ultramafic

lavas were produce d by deepe r mantle meltin g source, the

geochemical characteristics belonging to the oceanic class.

Most of the rock suites in the belt have been metamorphosed

to greenschist facies. However, its eastern margin is found to

be in thrust contact with the higher amphibolite facies rocks

(700°C at 6-7 Kbr), and the southern part near Mysore consist

of predominantly ultramafic rocks metamorphosed to amphibolite

and granulite facies. The northern part of the belt near Gadag

is least metamorphosed . Irrespectiv e of the grade of

metamorphism or of inferred ages of the various stratigraphic

groups, the belt shows a remarkable structural homogeneity of 3

phases of deformation from N to S and E to W and a convexity

towards East. Both major and minor Fl folds are tight isoclinal

with shallow to steep plunges and subvertical to subhorizontal

axial planes. The variation in the attitude and orientation of

the Fl axes has been controlled by the F2 episode which has

coaxially folded both the subparallel bedding and the first

generation axial plane schistocity cleavage. Only at Fl hinges

the intersection between SI and S2 is discernible. F3 is found

as general warps on F2 limbs. The Fl axial plane schistocity

cleavag e and F2 crenulatio n charge are generall y dippin g

(horizontal to subvertical) towards the east. High grade rocks

on the eastern margin have been thrust westwards over the low

grade central part. Structural data indicate considerable

crustal shortening along the belt. Inversion of stratigraphic

sequence is reflected, at many places by the youning directions

obtaine d from curren t bedding , grade d beddin g an d pillo w

couvexities. Horizontal compression and collision tectonics

therefore, appear to have played a significant role in the

development of the structural configuration of the belt.

As the 3000 m.y. old gray banded gneisses, found on the

eastern and western sides of the Chitradurga schist belt are

similar, the existing observations suggest the following two

possible models: i) The belt developed in a rift on the juvenile

Archaean continental crust which collapsed upon loading by

sediments, resulting in a shallow subduction and horizontal

compression. (ii) The belt evolved on an "Oceanic" crust between

two juvenile continental blocks to the East and West. Shallow

subduction and horizontal movemen t of the Eastern block would

then result in the present structural geometr y and consequent

welding of the two along this probable suture.
86

GEOCHEMICAL AND ISOTOPIC CONSTRAINTS ON THE TECTONIC SETTING OF THE

SERRA DOS CARAJAS BELT, EASTERN PARA, BRAZIL; Wm. J. Olszewski, Jr., Dept.

of Earth Sciences, Univ. of New Hampshire, Durham, NH 03824; Allan K. Gibbs

and Karl R. Wirth Institute for the Study of the Continents, and Dept. Geol.

Sci., Cornell University, Ithaca, NY 14853

The lower part of the Serra dos Carajas belt (Fig 1) is the metavolcanic

and metasedimentary Grao Para Group (GPG) (1-6). The GPG is thought to

unconformably overlie the older (but undated) Xingu Complex, composed of

medium and high-grade gneisses and amphibolite and greenstone belts. The

Lower Metavolcanic Sequence of the Grao Para Group (LMS) is estimated to be

about 4-6 km thick, consisting of massive, vesicular, and porphyritic mafic

volcanic flows and agglomeratic breccias and about 10-155& massive,

flow-banded, brecciated, and tuffaceous prophyritic rhyolite (6). The LMS is

overlain by the extensive, 100-400 m thick, and high-grade banded iron

formations of the Carajas Formation, followed by an Upper Sequence (US) of 1-3

km of mixed volcanic and clastic and chemical sedimentary rocks. The

stratigraphy of the US is poorly known, but it is thought to contain some

quartz-rich arenites, suggesting mature continental provenance (6). Much

thicker quartz-rich sandstones and conglomerates overlie the Upper Sequence,

with unknown degree of conformity.

Petrographic, geochemical, and isotopic analyses of the bimodal

metavolcanics of the LMS show these to be basalts, basaltic andesites,

trachyandesites (shoshonitic), and rhyolites (6,8). Spilitic alteration is

locally apparent, but the coherence of alkali element ratios and

readily-altered trace element compositions suggests that most samples did not

undergo strong alteration. Good correlation between HREES, Ti, and magnesium

number in the mafic rocks demonstrate the effects of fractional

crystallization in the mafic rocks. LREES, Si, K, Rb, Cs, and Ba do not

correlate with magnesium number, suggesting that variable enrichments of these

elements (fig. 2) reflect variable contamination of the basaltic melts with

crustal material. Several contamination components must have been involved,

since these elements are only weakly correlated among themselves, and with U,

Th, Nb, and Ta. Rhyolite patterns show significant negative Eu anomalies.

Zircons from two quartz porphyritic rhyolites give an age of 2758 + 39 Ma

(7), the best estimate of the age of eruption of the LMS. Rb-Sr whole-rock

analyses of mafic rocks yield an isochron of 2687 + 54 Ma, similar within the

range of calculated errors of the zircon age. Thus the GPG's Late Archean age

is well established. The high initial Sr isotopic ratio 0.7057 for the mafic

rock isochron is significantly higher than values of CHUR (0.7012, 9) or

depleted mantle (0.7008, 10) for 2758 Ma. This indicates contamination . by

older continental crust. Sm-Nd results are too restricted in distribution to

yield a usable isochron.

e Sr vs. e Nd values (Fig. 3) show a cluster around e Sr +50 and e Nd

+3. These indicate that the magma was more likely derived from a depleted

source than from a CHUR-like source. The high e Sr values are probably either

the results of seawater interaction, leaving Nd isotope ratios intact; or

contamination with older, presumably mafic crust that had elevated Rb/Sr

ratios, but mantle-like Sm/Nd ratios. One rhyolite has similar e Nd and e Sr

values, suggesting derivation from similar sources by similar processes.

Three of the mafic samples have negative e Nd and positive e Sr values,

possibly indicating contamination by older granulitic and granitoid crust.

Note that the ranges of diversity in the e Sr and e Nd data can be seen in the

basalts alone: the isotopic variation does not correlate directly with silica

content. Diverse sources of contamination are indicated, and might be found
SERRA DOS CARAJAS*", BRAZIL

Olszewski, W.J., Jr., Gibbs, A.K., and Wirth, K.R.

87

in the diverse lithologies of the underlying Xingu Complex.

The geochemical data indicate that the GPG has many features in common

with ancient and modern volcanic suites erupted through continental crust.

The mafic r"ocks clearly differ from those of most Archean greenstone belts,

and modern MORE, IAB, and hot-spot basalts. The geological, geochemical, and

isotopic data are all consistent with depostion on continental crust,

presumably in a marine basin formed by crustal extension. The isotopic data

also suggest the existence of depleted mantle as a source for the parent

magmas of the GPG. The overall results suggest a tectonic environment,

igneous sources, and petrogenesis similar to many modern continental

extensional basins, in contrast to most Archean greenstone belts. The

Hammersley basin in Australia and the circum-Superior belts in Canada may be

suitable Archean and Proterozoic analogues, respectively.

ACKNOWLEDGMENTS - Research in the Carajas region has been supported by the U.S. National Science

Foundation (under Grant No. EAR-8410379 to A.K.G.), with the generous institutional cooperation

of DOCBGEO. Some of the reported work will be discussed in greater detail in K.R.W.'s MSc.

dissertation, now in progreaa.

REFERENCES - [l] Hirata, W.K. , Rigon, J.C., Kadekaru, K., Cordeiro, A.A.C. , and Meirelea, E.A.

(1982) Anais Simp. Geol. Amazonia I, 100-110. [2J Santos, J-.O.S., and S.O.C. Loguercio (1934)

in Schobbenhaus et al. (eds.) Geologia do Brasil, DNPM, Brasilia, p. 93-127. L3j Amaral, G.

(1984) in F.?.M. de Almeida and Y. Hasui, 0 Precambnano do Braail, Edgard Blucher, Sao Paulo,

p. 6-35. [4] Gibbs, A.K., Wirth, K.R. , and Hirata, W.K. , (1985) Field report on a visit to the

Grao Para Group, Serra dos Carajaa, Brazil (unpublished), 35 p. l5j Cordani, U.G. , Tasainari,

C.C.G., and Kawashita, K. (1984) Ciencias da Terra, 9:6-11. L6J Gibbs, A.K., Wirth,

Hirata, W.K. , and Olszewski, Wm. J. Jr. (1985 submitted) II Simposio Geol

K • it • p

Amazonia, Soc.

Bras. Geol. Nucl. Norte, Dec. 1-9, 1985, Belem. 17 p. [7] Wirth, K.R., Gibbs, A.K., and

Olszewski, Wm. J. Jr. (1985 submitted) II Simp. Geol. Amazonia, Soc. Bras. Geol. Nucl.

Norte,.Dec. 1-9, 1985, Belem. 12 p. [s] Lemos, V.P., and Villas, R.N. (1983) Revista Bras.

Geocien. 13-165-177. [9] DePaolo, D.J., and Wasserburg, G.J. (1979) Geochim. Cosmochim. Acta

43:6156-617. [10] McCulloch, M.T., and Chappell, B.W. (1982) Earth Planet. Sci. Lett.

58:51-64. [ll] Cox, K.G. (1983) in Hawkesworth, C.L., and Norry, M.J. (eds.) Continental

Basalts and Mantle Xenoliths, Shiva Publ. Ltd. p. 139-157. [l2j Thirlwall, M.F., and Jon-a,

N.W. (1983) in Hawkeaworth, C.L., and Norry, M.J. (eds.) Continental Basalts and Mantle

Xenoliths, Shiva Publ. Ltd. p. 186-208. 50'00'W

SAMPLE LOCATIONS

SAVANNAS

GRANITE (1.8 Ga)

SANDSTONE UNIT

IRON FORMATION

GRAO PARA GROUP

XINGU COMPLEX

6* GO'S

6' 20' S
gg SERRA DOS CARAJAS, BRAZIL

Olszewski, W.J. Jr., Gitbs, A.K., and Wirth, K.R.

05

C» Rb Ba Tb. U K Ta Sr P Hf U Ce Nd Sm Eu Tb Yb

basalts

Fig. 2: Incompatable element

diagrams for maficf rocks, normalized for model primitive mantle of Wood et al. 1979, and

Hf, Eu, and Yb interpolated

from chondrite data. Karoo

basalts have similar enrichment patterns.

Fig. 3: Sr - Nd diagram for

Grao Para Group metavolcanic

rocks.

C« Kb Ba Th U K Ta Sr P HI La Ce Nd Sm Eu Tb Yb Tl

Depleted

Mantle n

cr

Q

T = 2758 Ma

• Basalt

O Basaltic Andeslt*

A RhyollU

Seawater Alteration

or Older Basalts

-5

-40

\

\ Crystal

Contamination

5O 100 150 200
N86-23120 89

POLYPHASE THRUST TECTONICS IN THE BARBERTON GREENSTONE BELT.

I. Paris. 2 Passage du Chantier, Paris 75012, France and, as of March 1, 1986,

Dept. of Geology, University of Canterbury, Christ Church, New Zealand.

In the circa 3.5 by old Barberton greenstone belt, the supracrustal rocks

form a thick and strongly deformed thrust complex. Structural studies in the

southern part of the belt have shown thai 2 separate phases of over-thrusting

(DI and D?) successively dismembered the original stratigraphy. Thrust

nappes were subsequently refolded during later deformations (Do an(j p.).

This poster deals with the second thrusting event which, in the study region

appears to be dominant, and (unlike the earlier thrusting), affects the entire

supracrustal pile.

The supracrustal rocks form a predominantly NE/SW oriented, SE dipping

tectonic fan (the Dp fan) in which tectonic slices of ophiolitic-like rocks

are interleaved witn younger sedimentary sequences of the Oiepgezet and

Malalotcha Groups1 (Fig. 1). Two distinct levels of decollement can be

distinguished within this fan: (1) Within the ophiolitic sequence, usually below

the pillow lavas. These zones are delineated by strongly sheared serpentinite

lenses and talcose schists. Asbestos fiber is commonly developed in such

sheared lenses, as for example in the Havelock and the Msauli asbestos deposits.

(2) At the base of the Diepgezet Group, within ferruginous shales and banded

cherts. This upper decollement zone is not always obviously sheared, but it is

ubiquitously folded in a disharmonic manner and is thought to have been gravity

induced, on a dynamic slope, during sedimentation, because: (1) The finely

laminated rocks at this stratigraphic level are conformably to unconformably

overlain by a 2 to 3 km thick medium to coarse grained clastic sequence (the

rest of the Diepgezet Group and the Malalotcha Group; the Malalotcha Group is

derived from a quartz-rich source and from the reworking of folded Diepgezet

Material). (2) Within the D2 fan, individual tectonic units may be folded

independently of one another (Fig. 2). The Dp foids are mostly isoclinal,

with fold axes broadly parallel to the thrust contacts (Fig. 2), and are

contemporaneous with the emplacement of the nappes. Another set of Do folds

is contemporaneous with the deposition of the Malalotcha Group sediments and

probably formed in tectonically ponded basins, during periods of thrust

propagation along the lower decollement level.

Structural and sedimentological data indicate that the Do tectonic fan

was formed during a prolonged, multi-stage regional horizontal shortening event

during which several types of internal deformation mechanisms were successively

and/or simultaneously active. Movement appears to have been predominantly to

the NW and to the N. During Dp, periods of quiescence and sedimentation

followed periods of thrust propagation. Although the exact kimematics which led

to the formation of this fan is not yet known, paleoenvironmental

interpretations together with structural data suggest that Dp was probably

related to (an) Archean collision(s).

References

(1) Lamb, S. (1984) Structures and sedimentology on the eastern margin of the

Archean Barberton greenstone belt, northwest Swaziland, Ph.D Thesis, University

of Cambridge, UK. (2) Paris, I.A. (1985) The geology of the Simbubule area,

Barberton. Ph.D Thesis, University of the Witwatersrand, Johannesburg, South

Africa.
90

THRUST TECTONICS IN THE BARBERTON BELT

Paris, I.

ORIGINAL PAGE IS

OF POOR QUALITY

FIG1

I00»

NW

i^JmS^iir'^

UfGEND

FIG 2.

CD H*«w/«» f"** CD DHtctomM

CD Leva | TKtonic unrt

CZ3 Clwrt ^ Ybwnang Olrtc

CD Pronimol focm iphvrotte V Ttewt

CD Hoo^Q.noco fcbic Mleauci **• 02 «««*»

.CD Oiitai focal tphtmd* V. D| lyndiM

OBllChM

Full

Fbull

FIC2

Figure Caption: (1) Simplified geological map of part of the Investigated area

(for location see inset). (2) Three sections, as located on Fig. 1 showing part

of the thrust complex. Note how some of the thrust-slices (individually

numbered) composed of sediments (Oiepgezet and Malalotcha Groups) are tight to

isoclinally folded. Folding and thrusting are related to the same regional

deformation (D-).
N86-23121

91

AGE CONSTRAINTS ON THE EVOLUTION OF THE QUETICO BELT,

SUPERIOR PROVINCE, ONTARIO; 3.A. Percival and R.W. Sullivan, Geological Survey

of Canada, 601 Booth Street, Ottawa, Ontario K1A OE8

Much attention has been focussed on the nature of Archean tectonic processes

and the extent to which they were different from modern rigid-plate tectonics. The

Archean Superior Province (1) has linear 'metavolcanic and metasediment-dominated

subprovinces of similar scale to Cenozoic island arc-trench systems of the western

Pacific (2), suggesting an origin by accreting arcs (3,4). Models for the evolution of

metavolcanic belts in parts of the Superior Province suggest an arc setting (4,5) but

the tectonic environment and evolution of the intervening metasedimentary belts are

poorly understood. In addition to explaining the setting giving rise to a linear

sedimentary basin, models must account for subsequent shortening and hightemperature, low-pressure metamorphism (6-8). Correlation of rock units and events

in adjacent metavolcanic and metasedimentary belts is a first step toward

understanding large-scale crustal interaction. To this end, zircon geochronology has

been applied to metavolcanic belts of the western Superior Province (9-13); this study

reports new age data for the Quetico metasedimentary belt, permitting correlation

with the adjacent Wabigoon and Wawa metavolcanic subprovinces.

The 10-100 km-wide Quetico belt extends at least 1200 km from beneath cover

in the west to the Kapuskasing structure and probably continues 800 km further east,

as the Opatica belt. It is mainly fault-bounded against adjacent metavolcanic rocks

but stratigraphic contacts are present locally. The belt consists of marginal zones of

metasedimentary schist and an interior zone of migmatite and granite. Marginal

metasediments have preserved sedimentary structures suggesting a homogeneous

sequence of turbiditic greywacke, possibly derived from adjacent volcanic highlands

(14). Conglomerate and cross-bedded sandstone of the Seine Group (15) occur

sporadically along the northern margin of the belt and have been interpreted as

proximal fan deposits of the Quetico turbidites (16) or as a younger sequence (15,17).

The most prominent structural features of the belt are the regular east-trending

bedding which dips steeply near the margins and moderately in the interior, and a

pervasive, gently east-plunging lineation. Several early sets of folds have been

recognized in detailed studies (18-20). Symmetrical low-pressure metamorphic

zonation characterizes marginal schists, where grade increases from chloritemuscovite at the margins, through biotite, staurolite, and garnet-andalusite zones, to

garnet-cordierite-sillimanite grade adjacent to the interior zone of migmatite and

intrusive granite. Common assemblages of garnet-andalusite throughout marginal

schists and locally in the interior indicate low metamorphic pressure (bathozone 2;

3.3 kbar (21)). Granulite facies occurs in the east near Flanders Lake (22) and

adjacent to the Kapuskasing zone (23), where metamorphic pressure is 4-6 kbar (24).

The regional metamorphic culmination is coincident with interior plutons, suggesting

that the granites transmitted heat to high levels in the crust.

Plutonic rocks, classified into three compositional groups, have restricted spatial

distribution: 1) a suite of small diorite-monzonite plugs cuts marginal schists and

extends locally into adjacent metavolcanic belts; 2) biotite-magnetite leucogranite

with local tonalite and amphibolite inclusions, occurs near the schist-migmatite

contact; and 3) peraluminous granite, with garnet, cordierite, muscovite, sillimanite,

apatite and tourmaline, are prevalent in the interior zone, particularly the Sturgeon

Lake batholith (8). Late pegmatites are ubiquitous in the interior zone and common in

the higher-grade parts of the marginal schist unit.
92

QUETICO BELT, SUPERIOR PROVINCE

Percival, 3.A. and Sullivan, R.W.

U-Pb zircon geochronology in the Wawa subprovince indicates major volcanic

activity between 27*9 and 2696 Ma (25) followed by DI deformation at about 2696,

deposition of alkaline ("Timiskaming") volcanics at 2689, D2 deformation, and intrusion

of post-tectonic plutons at 268* Ma (9) to 2668 Ma (26) (Fig. 1). In the Wabigoon

subprovince, volcanics were erupted in the interval 2755-2702 Ma, with post-tectonic

plutons younger than 2695 Ma (12) (Fig. l)i

A chilled porphyritic dacite sill cutting biotite-grade Quetico metasediments

yielded an imprecise U-Pb zircon date of 27*3 ±16 Ma, providing a minimum age for

sediment deposition. A single tonalite clast from metaconglomerate at Max Creek,

interpreted to be Seine equivalent, has zircons dated at 268* ± 10 Ma, interpreted as

the age of the source pluton. Together these dates show that the Quetico

metasediments and Seine Group are not facies equivalent. Monazites from the

geologically oldest plutonic rock type, a foliated biotite granite with zircons with

relict cores, are discordant, with an upper intercept of 268* Ma. Monazite from

massive peraluminous granite with probable inherited zircon is concordant at 2670 Ma.

Zircon and monazite from a pegmatite dyke form a discordia line with an upper

intercept of 2671 Ma (Fig. 1). The data do not permit definition of the length of time

of sediment deposition nor is the thickness of the sequence known; thus inferences on

iithospheric thickness (28) cannot be made.

Preliminary synthesis suggests that sediment deposition on extending crust

forming the Quetico basin probably occurred during volcanism in adjacent terranes,

possibly continuing until volcanism ceased. Closure of the basin during Dj and/or D2

events, dated in adjacent belts, led to folding of the sedimentary pile and thickening of

the weak crust. Conglomerates were deposited adjacent to marginal transcurrent

faults. During subsequent thermal relaxation, partial melts were extracted from lower

crustal metasedimentary and tonalitic rocks in a crustal root zone as well as from the

mantle. The derived granites and diorites ascended passively to within 10 km of the

surface, producing a regional low-pressure aureole in the host schists. A back-arc or

inter-arc setting is favoured over an accretionary prism environment for the Quetico

sediments because of its symmetry and high-temperature metamorphism which

probably occurred in a region of high heat flow.

REFERENCES

1) Goodwin, A.M. et al. (1972) Geol. Assoc. Can. Sp. Pap. 11, 527-623.

2) Hamilton, W. (1979) U.S. Geol. Surv. Prof. Pap. 1078, 3*5 p.

3) Langford, F.F. and Morin, 3.A. (1976) Am. 3. Sci. 276, 1023-103*.

*) Blackburn, C.E. (1980) Geoscience Canada 7, 6*-72.

5) Blackburn, C.E. et al. (1985) Geol. Assoc. Can. Sp. Pap. 28, 89-116.

6) Pirie, 3. and Mackasey, W.D. (1978) Geol. Surv. Can. Pap. 78-10, 37-*8.

7) Thurston, P.C. and Breaks, F.W. (1978) ibid, *9-62.

8) Percival, 3.A. et al. (1985) Geol. Surv. Can. Pap. 85-1A, 385-397.

9) Corfu, F. and Stott, G.M. (1985) Inst. Lake Superior Geol. 32 (abstr.), 15-16.

10) Turek, A. et al. (198*) Can. 3. Earth Sci. 21, *57-*6*.

11) Davis, D.W. and Edwards, G. (1982) ibid 19, 1235-12*5.

12) Davis, D.W. et al. (1982) ibid 19, 25*-266.

13) Nunes, P.O. and Thurston, P.C. (1980) ibid 17, 710-721.

1*) Ojakangas, R.W. (1985) Geol. Assoc. Can. Sp. Pap. 28, 23-*7.

15) Lawson, A.C. (1913) Geol. Surv. Can. Mem. *0, 111 p.

16) Wood, 3. (1980) Precambrian Res. 12, 227-255.

17) Poulsen, K.H. et al. (1980) Can. 3. Earth Sci. 17, 1358-1369.
',*'.. ' ; ;,; .QUETICO..BELT, SUPERIOR PROVINCE

PerciSral,'3.A. and Sullivan, R.W.

18) Kehlenbeck, M.M. (198*) Geoscience Can. 11. 23-32.

19) Borradaile, G.3. (1982) Precamb. Res. 19, 179-189.

20) Sawyer, E.W. (1983) ibid 22, 271-29*.

21) Carmichael, D.M. (1978) Am. 3. Sci. 278, 769-797.

22) Coates, M.E. (1968) Ont. Dep. Mines Geol. Rep. 68, 22 p.

23) Percival, 3.A. (1985) Geol. Surv. Can. Pap. 85-1A, 1-5.

2*) Percival, 3.A. (1985) Geol. Assoc. Can. Abstr. Prog. 10, A*8.

25) Turek, A. et al. (1982) Can. 3. Earth Sci. 19, 1608-1626.

26) Krogh, T.E. and Turek, A. (1982) ibid 19, 859-867.

27) Davis, D.W. (1985) Personal communication.

28) McKenzie, D. et al. (1980) Earth Planet. Sci. Lett. *8, 35-* 1.

ORIGINAL PAGE !S

OF POOR QUALITY 93

2680-

Wabigoon

Fig. 1: Age summary and tentative correlation diagram for the Wawa, Quetico and

Wabigoon subprovinces. Arrows crossing subprovince boundaries indicate sedimentary

provenance.
94 M86-23122

ARCHEAN MEGACRYSTIC PLAGIOCLASE UNITS AND THE TECTONIC SETTING OF GREENSTONES

W. C. Phinney, D. A. Morrison, NASA Johnson Space Center; D. Maczuga, LEMSCO,

Houston, TX 77058.

Large (up to 20 cm), equidimensional, commonly euhedral, plagioclase megacrysts of highly calcic composition (An80_Q0) occur commonly in all Archean

cratons in one or more of three distinct associations:

1) as cumulate crystal segregations of anorthosite or as megacrysts in basaltic dikes, sills, and flows in greenstone belts that vary in metamorphic grade

from greenschist to granulite. Throughout 100's of thousands of square kilometers of northwestern Ontario and Manitoba the plagioclase megacrysts occur

in pillowed and massive flows, sills, dikes, large inclusions in dikes, and

intrusive anorthositic complexes (Fig. 1) with areas of up to a few 100 km

and spanning a period of at least 100 m.y. in the 2.7 to 2.8 b.y. time frame,

2) as basaltic dike swarms in stable cratonic areas forming parallel to subparallel patterns over hundreds of thousands of square kilometers intruding

both granitic gneisses and supracrustal belts including greenstones. These

swarms include the Ameralik-Saglek system at 3.1 to 3.4 b.y. (Fig. 2) [1], the

Matachewan system at 2.5 to 2.6 b.y. [2], and the Beartooth-Bighorn system at

2.2 to 2.3 b.y. [3], and

3) as anorthositic complexes associated with marbles and quartzites (Sittampundi, India and Messina, South Africa) in granulite grade terrains.

Initial attempts to correlate tectonic settings of similar modern crystbearing units with their Archean counterparts were only partially successful.

Plagioclase phenocrysts of AngQ_gQ occur in basaltic volcanic flows in oceanic

crust at spreading ridges, hots'pots, aseismic ridges, and fracture zones [4],

These recent occurrences, however, normally involve only small phenocrysts up

to a few millimeters in size and usually more lathy than equidimensional in

shape [5], In contrast to these normal occurrences, volcanic flows over the

Galapagos hotspot display more equidimensional crysts up to 3 cm across [4].

Although these oceanic environments might be satisfactory tectonic analogs for

many greenstone occurrences, they certainly are not satisfactory for the extensive dike swarms in stable cratonic masses. Thus we turn for clues to a more

detailed understanding of the petrogenesis of the crysts and related melts.

The crysts are quite homogeneous, varying by little more than one to two

An units over several centimeters thereby suggesting nearly isothermal crystallization at nearly constant melt composition over the time required to grow

crystals commonly 6 to 8 cm across and up to 20 cm across and accumulate them

in large masses. Thin, more sodic rims on the order of 100 to 200 Aim wide are

common on large crysts when the groundmass plagioclase laths are more sodic

than the large crysts. The rims normally approach the composition of the

plagioclase in the groundmass (Table 1).

The nature of the parent melts, or melts in equilibrium with the large

crysts, has been an open question because: 1) the anorthositic complexes are

clearly cumulates with bulk compositions too rich in A^O^ and CaO to represent melts [6], and 2) the disparity in composition between plagioclase crysts

and plagioclase of the matrix suggests a lack of equilibrium between crysts

and the melt represented by their matrix.

Initial attempts to determine melt compositions by use of REE concentrations in megacrysts in conjunction with distribution coefficients for plagioclase and basaltic melts were fraught with problems resulting from modification of plagioclase REE concentrations by alteration, recrystallization, and

tiny inclusions. By utilizing several splits from each cryst in several

samples from the BVL anorthosite, mixing lines were determined and the least
Plagioclase Crysts and Greenstones

Phinney, W. C., Morrison, D. A and Maczuga, D. 95

modified REE concentrations were calculated for pristine plagioclases [7],

These values in conjunction with the most recent distribution coefficients

indicate melts with nearly flat REE patterns at 10X to 20X chondrites with

perhaps a slight depletion in the light REE's. The calculated patterns compare well with several cryst-bearing basalts in greenstone belts (Fig. 3) as

well as with the non-cryst-bearing basalts. These patterns are those of the

least enriched tholeiitic basalts which are very common in greenstone belts.

Comparison of these basalts with those in the cratonic dike swarms shows many

similarities (Fig. 3, Table 2) but the initial data suggests that the cratonic

dikes are slightly enriched in Si02, K2°»

 and light REE. It is tempting to

attribute these differences to contamination of the melts as they rise through

continental crust but the melts of the Galapagos when compared with MORB show

some of the same enrichments (Table 2) which in this case cannot be attributed

to continental contaminants. Further work on the pristine REE contents of

plagioclase megacrysts is underway and should help determine whether megacrysts in enriched melts formed from the more enriched or less enriched tholeiitic melts, or both.

At present the petrogenetic data require, at a minimum, isothermal crysplagioclase megacrysts from tholeiitic melts (the least engreenstone belts) followed by segregation of the plagioclase

then become entrained in rising melts to form intrusions or

Furthermore, the occurrences seem to require large volumes of

temperatures for long periods of time over huge areas having

tallization of

riched ones in

crystals which

volcanic flows,

melt at similar

both oceanic and cratonic associations. Continual generation of similar melt

and continuous addition of the melt to extensive networks of crystallizing

chambers is also strongly implied. The major remaining questions with significant implications for the setting and evolution of greenstone belts are: 1)

Does the cryst-producing melt have the same composition and crystallize under

the same conditions beneath greenstone belts, stable cratons, and current

oceanic crust? 2) Where do the plagioclase crysts form and accumulate; in low

or high pressure environments? 3) Is there a systematic change in the time of

megacryst emplacement across large areas such as might be produced by plates

overriding zones of melt production or other such time-dependent mechanisms?

Refs: [1] McGregor V. R. et al . (1985) Workshop: The Worlds Oldest Rocks, Lunar and Planetary

Inst., Houston. [2] Ernst R. £. (1982) Ont. Geol. Surv. Misc. Pap. 106, p. 53-56. [3] Miller J. D.

(1980-81) Wyo. Geol. Assoc. Earth Sc1. Bull., v. 13-14, p. 187-215. [4] Cullen A. et al. (1985)

preprint from Center for Vole. Univ. Oreg. [5] Blanchard D. P. et al. (1976) Jour. Geophys. Res.,

81. p. 4231-4246 and Donaldson C. (1977) EPSL, 37, p. 81-89. [6] Phinney W. C. (1982) LPI Tech.

Rpt. 82-01, Lunar and Planetary Inst., Houston, p. 121-124. [7] Morrison D. A. et al. (1985) Lunar

and Planet. Sc1. Conf. XVI, Lunar and Planetary Inst., Houston, p. 589-590.

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Plagioclase Crysts and Greenstones

Phinney, W. C., Morrison, D. A. and Maczuga, D.

PAGE

 'S

QUALITY

BIRD RIVER N. MATRICES UTIK LfiKE MfiTRICES

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N86-23123 97

RAINY LAKE WRENCH ZONE: AN EXAMPLE OF AN ARCHEAN SUBPROVINCE

BOUNDARY IN NORTHWESTERN ONTARIO; K. Howard Poulsen, Economic Geology and

Mineraloay Division, Geological Survey of Canada

Fig. li- Schematic diagram

illustrating structural

features of Rainy Lake

Wrench Zone. Short solid

arrows identify downward

facing units.

WABIGOON SUBPROVINCE

The Superior Province of the Canadian Shield comprises an alternation of

subrrovinces with contrasting litholoqical, structural and metamorphic

stylos (1). Rocks of the Rainy Lake area fornra fault bounded wedge between

tv;o of these subprovinces, the Wabigoon granite-greenstone terrain to the

north and the Quetico metasedimentary terrain to the south (Fig. 1). The

Ouetico and Seine River-Rainy Lake Faults bound this wedge within which interpretation of the stratigraphy has been historically contentious. In the

eastern part of the wedge, volcanic rocks and coeval tonalitic sills are unconformably overlain by fluviatile conglomerate and arenite of the Seine Group;

in the western part of the wedge, metamorphosed wacke and mudstone of the

Coutchiching Group are cut by granodioritic plutons. The Coutchiching Group

has previously been correlated with the Seine Group and with the turbiditic

Quetico metasediments of the Quetico Subprovince and these correlations are

the cornerstone of earlier tectonic models which relate the subprovinces (2,3).

The structural geology of the Rainy Lake area is characterized by the

following attributes:

(i) lenticular lithostratigraphic domains with discordant boundaries,

(ii) steep boundary faults,

(iii) regular orientation and sense of displacement of small ductile shear

zones,

(iv) regionally developed sub-vertical foliation which transects large lithological folds,

(v) shallow bimodal orientations of minor folds and lineations and a preponderance of folds of dextral asymmetry,

(vi) downward facing folds in the Rice Bay, Nickel Lake and Bear Pass areas

(arrowed, Fig. 1).

These observations compare favourably with the known characteristics of dextral

wrench or "transpressive" zones based both on experimental data and natural

examples (4,5,6,7,8). Much of this deformation involved the Seine Group, the

youngest stratigraphic unit in the area (9), and predates the emplacement of

late-to-post-tectonic granodioritic plutons for which radiometric data indicate

a Late Archean age.

The interpretation of a wrench zone separating the Wabigoon and Quetico

Subprovinces has important implications regarding the tectonic models which can
v- '\~ ,' *, VRAINY^LAKE; WRENCH ZONE

98 " K. Howard Poulsen

be used to relate them. Of great importance is the high probability that this

zone contains rocks which are actually allochthonous relative to those adjacent

in the Quetico and Wabigoon. Given this type of structural environment, not

only is correlation of stratigraphic Units between individual lenticular

domains difficult to establish simply on the basis of some lithological similarity but more important, the correlation with units exterior to the wrench

zone is even more suspect. Newgeochronological data (9) which demonstrates a

40 Ma difference in age between the Seine and Coutchiching strongly supports

this argument. Therefore the concept that Seine-type alluvial-fluvial rocks,

which are restricted spatially to the wrench zone are transitial "facies"

between Wabigoon volcanics and Quetico turbidites (2,3) finds little support

in a wrench zone interpretation.

PettiJohn (10) was the first to emphasize that Seine-type sedimentary

sequences occur all along the subprovince margin. Because these rocks also

correlate spatially with a well defined wrench zone it is instructive to inquire whether an alternate hypothesis might account for these observed rela-~

tionships without relying on the concept of facies equivalence. The link

between alluvial-fluvial sedimentation and wrench zones is well-known in

Cenozoic environments where thick alluvial, fluvial and lacustrine sequences

are restricted to narrow "pull-apart" basins associated with large transcurrent

faults (11,12,13). Such basins are localized by bends in marginal faults and

by intersections with fault splays. Lateral and vertical facies variations

are present within such basins (14) but these rocks are not contiguous with

rocks external to the basin. The size and geometry of the wrench fault system

at the southern margin of Wabigoon subprovince and the area! extent of the

Seine-type rocks are comparable with younger examples in which there is also

a juxtaposition of rocks of differing lithology. In many of these examples,

and possibly in the present one as well, the juxtaposed terranes have depositional histories which are quite independent so that present geographic geometry has no simple paleographic significance.

The proposal that wrench faulting is significant at the subprovince

boundary is not a new one. Hawley (15) first suggested a model of this type

for rocks in the Atikokan area to the east of Rainy Lake but the emphasis in

the past has been placed only on the late-stage displacements on the Quetico

Fault (2,16) rather than the possibility presented here that such late faulting

is merely a reflection of a broader zone of wrenching which also became a locus

for sedimentation.

References

(1) Stockwell, C. H. (1964) Geol. Surv. Can. Paper 65-17, 29p.

(2) MacKasey, W. D., Blackburn, C. E. and Trowell, N.F. (1974), Ont. Div.

Mines. MP 58. 30p.

(3) Goodwin, A. M. (1977~TC~an. J_, Earth Sci., 1_4, p. 2737-2759.

(4) Wilcox, R. E., HardingTT. P. and Seely, D. R. (1973) AA7G Bull., 57,

p. 74-96.

(5) Tchalenko, J. S. (1970) Geol. Soc. Am. Bull.. 81_, p. 1625-1640.

(6) Lowell, J. D. (1972) Geol. SocT~Am. Bull., 83_, p. 3091-3102.

(7) Freund, R. (1970) GeoTTS'ocT'A'm. Bull., 78, p. 188-200.

(8) Ramsay, J. G. (198UT~J. Struct. Geol.. 2, p. 83-99.

(9) Davis, D. W. and Corfu, F. (this~voTumeT

(10) Pettijohn, F. J. (1943) Geol. Soc. Am. Bull.. 54_, p. 925-972.
RAINY LAKE WRENCH ZONE

K. Howard Poulsen 99

(11) Schubert, C. (1980) vK Struct. Geol., 2, p. 463-468.

(12) Morris, R. J. and Carter. R. M."TT9"82)"Tectonophys1cs, 87, p. 11-24.

(13) Miall, A. D. (ed.) (1981) Geol. Assoc. Can., Sp. Paper 23". Z7?y~.

(14) Mann, P. and Burke, K. (19%2T"Gefcl. Soc7"Aln. Abst. Ws , New~0"rleans, p. 555.

(15) Hawley, O.E. (1930) J.. Geol., 387"p.~57l-5T7.

(16) Schwerdtner, W. M., Stone, D. ,~0~sadetz, K., Morgan, J. and Stott, G. M.,

(1979) Can. J. Earth Sci.,,16, p. 1965-1977.
N86-23124

A PALAEOMAGNETIC PERSPECTIVE OF PRECAMBRIAN TECTONIC STYLES

P,W« Schmidt and B.J.J. Embleton, CSIRO .Division of Mineral Physics and

Mineralogy, P.O. Box 136, North Ryde NSW Australia

'<V

The considerable success derived from palaeomagnetic studies of

Phanerozoic rocks with 'respect to the tectonic styles of continental drift (1)

and plate tectonics (2), etc. have not been repeated by the many

palaeomagnetic studies of Precambrian rocks. This is undoubtedly related to

the vast amount of Precambrian time compared with Phanerozoic time, and the

concomitant uncertainties of magnetisation ages and rock ages, yet it is still

surprising that there is little evidence of consolidation or even convergence

of opinions regarding tectonic styles prevalent during the Precambrian. After

all, there are 30 years of research with results covering the major continents

for Precambrian times that overlap considerably yet there is no concensus even

in the grossest terms. There is good evidence that the usual assumptions

employed by palaeomagnetism are valid for the Precambrian which only serves to

exacerbate the problem. The existence of magnetic reversals during the

Precambrian, for instance, is difficult to explain except in terms of a

geomagnetic field that was predominantly dipolar in nature. It is a small

concession to extend this notion of the Precambrian geomagnetic field to

Include its alignment with the Earth's spin axis and the other virtues of an

axial geocentric dipole that characterise the recent geomagnetic field. In

addition it is not a forceful argument to claim that early studies of

Precambrian rocks need to be re-done, since re-studies have often only served

to confirm the early works. Therefore we submit that the palaeomagnetic

results derived from Precambrian rock units are not easily dismissed. It is

simply untenable that the majority of the data are spurious and claims that

synopses of Precambrian data are invalid, cannot be sustained in such terms.

Such arguements posed against the evidence for continental drift have long

been debunked. There are, nevertheless, differing interpretations of

Precambrian palaeomagnetic data and it is the purpose of this brief article to

address this problem.

Methods that have been used to interpret Precambrian palaeomagnetic data

fall into two classes. The first class assumes the existence of a "Pangaea"

or some supercontinent and proceeds to use the palaeomagnetic data, a_

posteriori, to support the model. The second class, which we prefer, accepts

the palaeomagnetic data at face value (as synthesised by workers closely in

touch with the results) and proceeds to view the overall relationships of the

data, isolated from preconceived notions. This latter approach has led us to

suggest that the present day geographical relationships of continents (from

which a reasonable amount of data for the Precambrian are available) yields

the more satisfactory comparison. Of course small adjustments of the

continents refine this comparison, but overall an excellent agreement in

Precambrian pole paths can be realised by leaving the continents in their

present locations.

Limitations of the available data in our earlier comparisons (3)

restricted the time span of comparisons between different continents to 2300

Ma - 1900 Ma for North America and Africa and 1800 Ma - 1600 Ma for North

America, Greenland and Australia. Recently two results have been derived from

igeneous rock about 2900 Ma in age, in Australia and Africa. The

palaeomagnetic pole positions from these rock units are in close proximity,
PERSPECTIVE

Schmidt, P.W. and Embleton, B.J.J. 101

suggesting that the present geographic relationship of Australia and Africa is

valid for 2900 Ma ago. The pole position from the Millindinna Complex,

Australia, dated at 2860±20 Ma is at 11.9°S, 161.3°E, dp=6.8°, dm=8.4°(4),

while the pole position from the Usushvana Complex, Africa, dated at 2880 Ma

is at 11.6°S, 165.8°, dp=-5.1°, dm=7.5°(5). Thus there is evidence that during

the Precambrian North America and Australia were in their present relative

geographic locations for 1800 Ma-1600 Ma, as were North America and Africa for

2300 Ma-1900 Ma, and now Africa and Australia, at least for 2900 Ma ago.

These observations are not easily reconciled with Phanerozolc

palaeomagnetic results as we have already discussed(3), but they are a matter

of record and must be explicable. In terms of greenstone terranes it is

obvious that tectonic models postulated to explain these observations are

paramount in understanding Precambrian geology. What relevance the current

geographical relationships of continents have with their Precambrian

relationships remains a paradox, but It would seem that the ensialic model for

the development of greenstone terranes is favoured by the Precambrian

palaeomagnetic data.

(1) Irving, E., 1964. Paleomagnetism and its application to geological and

geophysical prpblems. Wiley, N.Y., pp. 399.

(2) McElhinny, M.W., 1973. Palaeomagnetism and plate tectonics. Cambridge

University Press, Cambridge, pp. 358.

(3) Embleton, B.J.J. and Schmidt, P.W., 1979. Nature ^82_, 705-707.

(4) Schmidt, P.W. and Embleton, B.J.J., 1985. J. Geophy. Res., 90, 2967-

2984.

(5) Layer, P.W., Kroner, A. and McWilliams, M.O., 1985, pers. comm.
102

H86-23125

THE WISCONSIN MAGMATIC TERRANE: AN EARLY PROTEROZOIC GREENSTONEGRANITE TERRANE FORMED BY PLATE TECTONIC PROCESSES; Klaus J. Schulz, U.S.

Geological Survey, Reston, VA 22092 and Gene L. LaBerge, Department of

Geology, University of Wisconsin-Oshkosh, Oshkosh, WS 54901 and U.S.

Geological Survey

The Wisconsin magmatic terrane (WMT) is an east trending belt of

dominantly volcanic-plutonic complexes of Early Proterozoic age ("18bO m.y.)

that lies to the south of the Archean rocks and Early Proterozoic epicratonic

sequence (Marquette Range Supergroup) in Michigan. It is separated from the

epicratonic Marquette Range Supergroup by the high-angle Niagara fault, is

bounded on the south, in central Wisconsin, by Archean gneisses, is truncated

on the west by rocks of the Midcontinent rift system, and is intruded on the

east by the post-orogenic Wolf River batholith.

Although the history of the WMT is complex in detail, integration of

recent studies (Sims and others, in press) provides an overview of its nature

and evolution. The WMT shows many similarities to Archean greenstone-granite

(AGG) terranes (Condie, 1981). In fact, until recent U/Pb zircon dating,

considerable controversy existed as to the age of the rocks of the WMT.

Insofar as the comparisons between the WMT and AGG terranes are valid, understanding of the tectonics of the WMT may provide important insights into the

tectonic processes involved in the evolution of at least some AGG terranes.

As in many AGG terranes, a major portion of the WMT is comprised of

volcanic rocks and lesser volcanogenic sediments variably metamorphosed to

lower greenschist to amphibolite facies. The supracrustal rocks show a

complex stratigraphy with at least three successions distinguished on the

basis of differences in composition, metamorphism, and structural fabric

(LaBerge and Myers, 1984; Sims and others, in press). The older units are

dominantly subaqueous basaltic lavas and consanguineous intrusive rocks

which are overlain locally by intermediate to felsic volcanic and volcaniclastic units, some in part subaerial (LaBerge and Myers, 1984). Both bimodal

(basalt-rhyolite) and calc-alkaline (basaltic andesite through rhyolite)

suites are present with the former hosting volcanogenic massive sulfide

deposits (May and Schmidt, 1982). Komatiites have not been recognized within

the WMT. The older basaltic units are dominantly tholeiitic in character,

show strong to moderate depletion of light REE elements ([La/Yb]N=0.09-0.89)

and high-field-strength elements (Hf, Zr, Ta, etc.), and are lithologically

and compositionally similar to recent back-arc basin basalts (e.g. Mariana

Trough, Wood and others, 1981), island-arc tholeiites (e.g. Scotia arc,

Hawkesworth and others, 1977), and some ophiolitic basalts (e.g. Troodos,

Kay and Senechal, 1976). The younger calc-alkali ne units are enriched in

LIL elements ([La/Yb]N=2.5-9.5), are also depleted in high-field-strength

elements, and are similar to volcanic sequences found in recent island-arcs

(e.g. Sunda Arc, Whitford and others, 1979).

Sedimentary rocks are locally found to overlie and(or) interfinger with

the volcanic succession. They include graywacke, argillite, thin ironformation, chert, and minor conglomerate, some containing granitoid boulders

(LaBerge and Myers, 1984).
Wisconsin Magmatic Terrane

103

Schulz, K. J. and LaBerge, fi. L.

Intrusive rocks within the WMT appear to have been largely diapirically

emplaced and show a temporal progression from gabbro and diorite through

tonalite and granite. They range from calcic to calc-alkaline in character,

although locally slightly alkaline varieties are also present (Sims and

others, 1985). The granitoids show an overall increase from north to south

across the terrane in their average I<20/Na20 ratios and Si02 contents.

Gneissic rocks, found in domes and block uplifts, are mostly tonalite to

granodiroite and are also calc-alkaline (Sims and others, 1985). Both

lithologically and chemically, the WMT granitoids appear similar to those

formed at compressional plate-margins (Brown, 1982).

Ultramafic rocks are present in the WMT, particularly along the northern

and southern margins. They are mostly serpentinized, but perioditic and

pyroxenitic lithologies are recognized. These ultramafic rocks are often

spatially associated with gabbroic rocks and were in some cases structurally

implaced. The ultramafic-gabbroic bodies are lithologically and chemically

similar to recent ophiolitic fragments.

Structure within the WMT is complex and consists regionally of large

structural blocks having diversely oriented internal structures that are

bounded by ductile deformation zones ("shear zones"; LaBerge and Myers, 1984;

Sims and others in press). Within the blocks, the supracrustal rocks show

generally steep dips and open to isoclinal folds. The deformation zones

bounding the blocks record pronounced flattening in the foliation planes

and a strong component of verticle movement (Palmer, 1980). This intense

deformation along zones is regional in scope, and generally younger than

the pervailing internal structural fabric within the blocks. Domes along

the northern margi n of the terrane, representing large-scale, antiformal

fold-interference structures, modified by diapirism and by intrusion of

granitoids, have further deformed and metamorphosed the mantling supracrustal

rocks (Sims and others, 1985).

U-Th-Pb zircon ages on the volcanic and associated gneissic and granitoid rocks that comprise the WMT (VanSchmus, 1980; Sims and others, in press)

indicate that they formed from 1,890 to 1,830 Ma. Detailed isotopic dating

in the northeastern portion of the WMT (Sims and others, 1985) indicates that

volcanism, granitoid intrusion, metamorphism, and deformation within this

region occurred from 1,865 to 1,835 Ma ago, a time span of 30 m.y.

The overall lithologic, geochemical, metallogenic, metamorphic, and

deformational characteristics of the WMT are similar to those observed in

recent volcanic arc terranes formed at sites of plate convergence. It is

concluded that the WMT represents an evolved oceanic island-arc terrane

accreated to the Superior craton in the Early Proterozoic. This conclusion

is strengthened by the apparent absence of Archean basement from most of the

WMT, and the recent recognition of the passive margin character of the

epicratonic Marquette Range Supergroup (Larue and Sloss, 1980). On the basis

of the new data for the WMT and the epicratonic sequence in Michigan, Schulz

and others (in press) have proposed the following tectonic model: 1) early

crustal rifting and spreading along the southern margin of the Superior

craton, 2) subsequent subduction and formation of a complex volcanic arc,

and, 3) with oblique convergence, collision of the arc with the continental

margin (epicratonic) sequence and Archean crust of upper Michigan culminating
Wisconsin Magmatic Terrane

104 Schulz, K. J. .and LaBerge, G. L.

in the Penokean orogeny. This tectonic model is similar to plate tectonic

histories recently presented for other Early Proterozoic terranes of North

America (Hoffman, 1980; Lewry, 1981; Karlstrom and others, 1983). This

indicates that the events and processes occurring in the Lake Superior region

were not unique, and that the tectonic processes operating were generally

similar to those recognized for the Phanerozoic. Given the general similarity

of some AGG terranes to the Early Proterozoic magmatic terranes, it seems

likely that subduction and plate collisions were also operative in the Archean.

References:

Brown, G. C., 1982, In: Thorpe, R. S., ed., Andesites; New York, John Wiley

and Sons, p. 437-461.

Condie, K. C., 1981, Archean Greenstone Belts: Elsevier Scientific Publ. Co.,

New York, 434 p.

Hawkesworth, C. J., O'Nions, R. K., Pankhurst, R. J., Hamilton, P. J., and

Evensen, N. M., 1977, Earth and Planet, Sci. Letters, v. 36, p. 253-

262.

Hoffman, P. F., 1980, In: Strangway, D. W., ed., Geol. Assoc. Canada Special

Paper 20, p. 523-549.

Karlstrom, K. E., Flurkey, A. J., and Houston, R. S., 1983, Geol. Soc.

America Bull., v. 94, p. 1257-1274.

Kay, R. W., and Senechal, R. G., 1976, Journal Geophy. Research, v. 81,

p. 964-970.

LaBerge, G. L., and Myers, P. E., 1984, Geol. Soc. America Bull., v. 95,

p. 246-253.

Larue, 0. K., and Sloss, L. L., 1980, Geol. Soc. America Bull., Part I, v. 91,

p. 450-452.

Lewry, J. F., 1981, Nature, v. 294, p. 69-72.

May, E. R., and Schmidt, P. G., 1982, Geol. Assoc. Canada Spec. Paper 25,

p. 447-480.

Palmer, E. A., 1980, Unpublished M.S. thesis, Univ. of Minnesota-Duluth,

Duluth, Mn, 127 p.

Schulz, K. J., Sims, P. K., Morey, G. B., in press, Precambrian: Conterminous

U.S., DNAG v. C-2, Geol. Soc. America.

Sims, P. K., Peter-man, Z. E., and Schulz, K. J., 1985, Geol. Soc. America

Bull., v. 96, p. 1101-1112.

Sims, P. K., Schulz, K. J., Peterman, Z. E., and VanSchmus, W. R., in press,

Precambrian: Conterminous U.S., ONAG v. C-2, Geol. Soc. America.

VanSchmus, W. R., 1980, In: Morey, G. B., and Hanson, G. N., eds., Geol.

Soc. America Special Paper 182, p. 159-168.

Whitford, D. J., Nicholls, I. A., and Taylor, S. R., 1979, Contrib. Mineral

Petrol., v. 70, p. 341-356.

Wood, D. A., Marsh, N. G., Tarney, J., Joron, J.-L., Fryer, P., and

Treuil, M., 1981, In: Hussong, D. M., and Uyeda, S., and others (eds.)

Init. Rep. Deep Sea Drilling Proj. 60, p. 611-645, U.S. Government

Printing Office, Washington, D. C.
N86-23126 105

NEW INSIGHTS INTO TYPICAL ARCHEAN STRUCTURES IN GREENSTONE TERRANES

OF WESTERN ONTARIO: W.M. Schwerdtner, Department of Geology, University of

Toronto, Toronto, Canada M5S 1A1

Ongoing detailed field work in selected granitoid complexes of the

western Wabigoon and Wawa Subprovinces, southern Canadian Shield, has led to

several new conclusions: (1) Prominent gneiss domes are composed of

prestrained tonalite-granodiorite and represent dense hoods of magmatic

granitoid diapirs. The diapiric material commonly was a syenite-diorite

crystal mush. (2) The deformation history of the prestrained gneiss remains

to be unraveled. (3) The gneiss lacked a thick cover of mafic metavolcanics

or other dense rocks at the time of magmatic diaprisim. (4) The synclinoral

structure of large greenstone belts is older than the late gneiss domes and

may have been initiated by volcano-tectonic processes. Multi-phase granitoid

plutonism greatly tightened the synclinoria. (5) Small greenstone masses

within the gneiss are complexly deformed, together with the gneiss. (6) No

compelling evidence has been found of ductile early thrusting in the gneiss

terranes. Zones of greenstone enclaves occure in hornblende-rich

contaminated tonalite and are apt to be deformed magmatic septa. Elsewhere,

the tonalite gneiss is biotite-rich and hornblende-poor.

These conclusions rest on several new pieces of structural evidence.

(1) Oval plutons of syenite-diorite have magmatic strain fabrics and sharp

contacts that are parallel to an axial-plane foliation in the surrounding

refolded gneiss. (2) Gneiss domes are lithologically composite and contain

large sheath-like structures which are deformed early plutons, distorted

earlier gneiss domes, or early ductile nappes produced by folding of planar

plutonic septa. (3) The predomal attitudes of gneissosity varied from point

to point. It is difficult to prove by conventional structural methods what

caused the state of early deformation in the large gneiss domes. New

approaches are being developed based on the patterns of total and incremental

finite strain in the granitoid terranes under study. 1
106

DEFORMATIONAL SEQUENCE OF A PORTION OF THE MICHIPICOTEN

GREENSTONE BELT, CHABANEL TOWNSHIP, ONTARIO; Catherine H. Shrady and

George E. McGill, Dept. of Geology and Geography, University of

Massachusetts, Amherst, MA 01003

Detailed mapping at a scale of one inch = 400 feet is being carried

out within a fume kill, having excellent exposure, located in the southwestern portion of the Michipicoten Greenstone Belt near Wawa, Ontario.

A simplified geological map of the area described here is presented

in a companion abstract (Fig. 1 in 1).

The rocks are metasediments and metavolcanics of lower greenschist

facies. U-Pb geochronology indicates that they are at least 2698 ± 11 Ma

old (2). The "lithologic packages" (1) strike northeast to northwest,

but the dominant strike is approximately east-west. Sedimentary structures and graded bedding are well preserved, aiding in the structural

interpretation of this multiply deformed area.

Deformation in this area is tentatively divided into six phases

(0-5). Phase 0 is soft sediment deformation. Folds of this type are

generally small (amplitudes ranging from several millimeters to tens of

centimeters); however, some early larger scale (up to 10 meters in

amplitude) tight to isoclinal folds with no or a very poorly developed

axial plane cleavage may be slump folds.

Included within Phase 1 of deformation is the regional overturning

resulting in rocks that dip north and young to the south in the northern

part of our area and extending well to the north (1,3,4), and rocks

that dip south and young north in the southern part of our area and

farther south (l,3a). To what extent the regional steep dips are attributable to this phase of deformation or to later refolding is, at present,

not known. Also included within Phase 1 are an approximately bedding

parallel cleavage, and pebbles within conglomeratic units flattened

parallel to this cleavage. It is thought that these latter two features

are associated and likely relate to the regional overturning.

Cut by and therefore pre-dating Phase 2 cleavage, but of uncertain

temporal relationship to the structures included within Phase 1, are

areally significant faults that separate lithologic packages. These

faults regionally follow but locally truncate bedding. In places,

they are associated with an apparently old fracture cleavage.

Phase 2 is characterized by a penetrative northwest to north striking cleavage of moderate dip. Phase 2 cleavage crenulates Phase 1

cleavage where both are clearly present; however, in much of the area,

these two cleavages cannot be separated. Related examples of mesoscopic

folds are rare, and associated structures of regional significance

have not been recognized.
> N86-23127

DEFORMATIONAL SEQUENCE, MICHIPICOTEN BELT

Shrady, C.H. and McGill, G.E. 107

Phase 3 cleavage is penetrative where well developed and crenulates

both Phase 1 and Phase 2 cleavages. Within the area mapped, Phase 3

cleavage strikes northeast with generally steep northwest or southeast

dips; dip direction and angle commonly change within individual outcrops.

Dips as low as 30° are locally present in the northwest part of the area.

It is not clear whether variation in dip indicates the existence of

two distinct northeast striking cleavages or whether it is due to later

minor folding about sub-horizontal axial surfaces. Phase 3 cleavage

is axial planar to folds that are open to tight, range in scale from

several millimeters to tens of meters in amplitude, and refold earlier

folds. At one locality, Phase 3 cleavage and associated folds appear

related to late movement on a fault that approximately parallels bedding.

It is not yet clear if this fault is entirely young, or whether it

is a reactivated older structure.

Steeply dipping northeast and north-northwest trending faults

constitute Phase 4. However, some movement on these faults post-dates

diabase dikes (Phase 5) that trend north-northwest and northeast.

Locally developed fracture cleavages appear to be associated with diabase

dike emplacement, but because the dikes commonly follow trends of older

faults, some or all of these fracture cleavages may be related to the

faults rather than to the dikes.

In summary: we have tentatively identified at least six phases of

deformation within a relatively small area of the Michipicoten Greenstone

Belt. These include the following structural features in approximate

order of occurrence: 0) soft-sediment structures; 1) regionally overturned rocks, flattened pebbles, bedding parallel cleavage, and early,

approximately bedding parallel faults; 2) northwest to north striking

cleavage; 3) northeast striking cleavage and associated folds, and

at least some late movement on approximately bedding parallel faults;

4) north-northwest and northeast trending faults; and 5) diabase dikes

and associated fracture cleavages. Minor displacement of the diabase

dikes occurs on faults that appear to be reactivated older structures.

References

(1) McGill, G.E. and Shrady, C.H. (1986) This volume.

(2) Turek, A., et al. (1984) Can. Jour. Earth Sci., 21, 457-464.

(3) Sage, R.P., et al. (1982) Ontario Geol. Surv. Prelim. Maps of

Precambrian Geology: a. Chabanel Twp., Map P. 2439; b. Esquega

Twp., Map P. 2440; C. McMurray Twp., Map P. 2441; d. Lastheels

Twp., Map P. 2442; e. Musquash Twp., Map P. 2568.

(4) Sage, R.P., et al. (1984) Prelim, map of Precambrian geology of

Corbiere Twp., Ontario Geol. Surv. Map P. 2607.
108

ORIGINAL PAGE IS

OF POOR QUALfTY N86-23128

A CONTINENTAL RIFT NOOEL FOR THE LA GRANDE 6REENSTONE BELT;

T.Skulski (1), A.Hynes (1), M.Liu (2), D.Francis (1), B.Rivard (1),

K.Stafflatelopoulou-Seynour (3).(1) Departnent of Geological Sciences, McGill

University, Montreal, Canada, <2) Department of Geosciences, University oi

Arizona, Tucson, Arizona, (3) Department of Geology, Concordia University,

Montreal , Canada.

Stratigraphic relationships and the geochemistry of volcanic rocks

constrain the nature and tiding of the tectonic and magmatic processes in the

pre-deformational history of the La Grande greenstone belt in the Superior

Province of north-central Quebec (Fig. 1). With the exception of a locality

in the western part of the belt the lowermost supracrustals in this belt are

obscured by syntectonic granitoid intrusives. The supracrustal succession in

the western part of the belt consists of a lower sequence of immature clastic

sediments and mafi c volcanoclastics, overlain by pillowed and massive basalts

(Fig. 1, A-A'). Further east, along tectonic strike, a lower sequence of

mafic volcanoclastics and immature clastic sediments is overlain by a thick

sequence of pillowed and massive basalts, and resedimented coarse clastic

sediments and banded iron formation. These are overlain by massive basaltic

andesites, andesites and intermediate volcanoclastics intercalated with

immature clastic sediments (Fig. 1, B-B'). In contrast, in the eastern part

of the belt lenses oi felsic volcanics and volcanoclastics occur at the base

of the succession and pillowed and massive basalts are overlain by komatntes

at the top (Fig. 1, C-C').

The

reflect

clastic

addition

volcani c

unconformable

succession and

lower sequences of clastic sediments in the central part of the belt

a mixed intrabasinal and extrabasinal provenance, but the upper

sediments have a uniquely extrabasinal tonalitic provenance. In

metasedimentary and granitoid xenoliths have been found in the

pil e in the central and eastern parts of the belt and a local

contact is believed to exist between the supracrustal

an underlying tonalitic basement in the west (1). Therefore a

model in which the La Grande belt formed on a sialic crust is favoured.

The largest volumes of

tholeiitic basalts (Fig. 2).

liquids, but have undergone a

eruptive rocks in the La Grande belt are

These basalts are not primary mantle-derived

polybaric fractionation history (1, 2 and 3).

Their parental magmas are

believed to have been

basaltic komatiites (Fig.

2). The basaltic komatiites

and most magnesian basalts

lie along a steep slope in

Al-Si space (Fig. 2) which

is best explained by the

fractional crystallization

of orthopyroxene and

olivine (4, 1). Coexistence of these two

silicate phases and a

liquid of basaltic

composition is restricted

Figure 1 Geology of the La Grande greenstone belt.
A CONTINENTAL RIFT MODEL FOR LA GRANDE

Skulski T., Hynes A., Liu M., Francis D. et al. 109

AIX

o

1 OlIV+OPX

20UV+SP

30UV+PLAG+CPX

30 OLIV Cj y OPX 60

60

1OLIV+-OPX

2OIIVVSP

3OllV-l-PLAG-fCPX

fez

Figure 2 Al-Si and Mg-Fe in cation?.. The

solid line encloses basalts from section

A-A', dotted line is basalts •from section

B-B', dash-bar and dash-dot are komatiites

and basalts respectively from section C-C

and the dashed line includes komatiites

and basalts from Lac Buyer (north of C-C).

of rhyolitic magmas in the east (4).

komatiitic magmas ingested sialic

crystallization and were ultimately

andesites. These contaminated magmas

to pressures on the order of 10 kb

(5). Thus the basalts represent

komatiiti c liquid s which have been

modified by differing extents of

fractionation at depths on the

order of 30 km before migrating to

higher levels in the crust (3, 1

and 6). A spectrum of basaltic

compositions are found in the La

Grande belt of which the

endmembers are an Fe-enriched

suite and those which have

negligibl e Fe variation (Fig. 2).

The Fe-enriched basalts have

undergone extensive low pressure

fractionation of a gabbroic

assemblage, which is probably the

result of a more protracted

residence time in upper crustal

conduit system than the relatively

constant Fe group. The degree of

fractionation of the komatiiti c

liquid s and their location in

space and time may reflect the

variable efficiency of a crustal

density filter (cf. 7). Thus, the

occurrence of komatiiti c lavas in

the upper levels of the

supracrustal succession may be due

to late failure of the crustal

barrier. Their restricion to the

eastern parts of the belt may

reflect development of a major

rift only there. Ponding of mafic

magmas withi n the sialic crust may

have resulted in the melting of

the crust and the early eruption

Toward the central parts of the belt,

crust, were modified by fractional

erupted as basaltic andesites and

are characterized by high compatible

18

element (eg. Ni and Cr) and fractionated, enriched light rare earth element

abundances (up to 100X chondrite) (8).

The La Grande greenstone belt can be explained as the product of

continental rifting (6). The restricted occurrence of komatiites, and

eastwardly directed paleocurrents in clastic sediments in the central part of

the belt are consistent with rifting commencing in the east and propagating

westward with time (Fig. 3). The increase in depth of emplacement and

deposition with time of the lower three units (Fig 1, section B-B') in the

central part of the belt reflects deposition in a subsiding basin (6). These

supracrustal rocks are believed to represent the initial rift succession

(c.f. 9). Model calculations (Fig.3) reveal that the extension factor for

lithosphere neccessary to account for the observed initia l subsidence in the
no

A CONTINENTAL RIFT MODEL FOR LA GRANDE

Skulski T., Hynes A., Liu M., Francis D. et al,

central part of the belt (6) is comparable in magnitude with that measured in

Modern sedimentary basins where the continental lithosphere is believed to

have been rapidly thinned (10). The occurrence of clastic sediments of

granitic provenance high in the succession in the central parts of the belt

may reflect the uplift and erosion of marginal forebulges that formed as a

result of lithospheric flexure.

"P Figure 3 Initia l elevation change

_SJ versus uniform extension factor. For

an initia l elevation change of .9 km

corresponding to the subsidence that

is observed in the lower three units

of section B-B'corrected for the

basin fil l and 1 km of water requires

a uniform extension factor of

approximatel y 1.5. The symbols used

are: crustal thickness (tc), crustal

and mantle densities (pc) and (pm)

respectively, temperature at the base

of the slab (T) and lithosphere

thickness (A).The thermal expansion

coefficient used is 3.2 x id C. The

calculations were performed using the

method of Royden and Keen (11).

Ill

0

z

i

o

o

1-

UJ

UJ

^,

O

-1

-3

4

i i i i i — i i

tc •

PC -

pm •

A -

k

Ki^^=

=

1 234567 8

— i — i • i- »

36 km

2.89 gm on"3

333 gm cm"3

80km

T-I394C* .

9 10 It 12 13

UNIFORM EXTENSION

FACTOR

Rivard B. and Francis D. (1986) Contrib. Mineral. Petrol, (in preparation).

Skulski T. , Hynes A. and Francis D. (1984) in Chibougamau-Stratigraphy and

Mineral ization, CIM special V_. 34, Guha J^ and Chown E_. editors, p. 57-72.

Stamatelopoulou-Seymour K. , Francis D. and Ludden J. (1983) Contrib. Mineral.

Petrol. . 84_, JK 6-14.

Stamatelopoulou-Seymour K. and Francis D. (1986) Precamb. Res. (in

preparation).

O'Hara M. (1968) Earth-Sci Rev., 4_, p. 69-133.

Skulski T. , Hynes A. and Liu M. (1986) Precaab. Res, (in preparation).

Francis D. , Ludden J. and Hynes A. (1983) J. Petrol. , 2_4, JK 556-582.

Skulski T. (1986) unpublished MSc thesis, Department p_f_ Geological Sciences,

McGill University.

Bickle M. and Eriksson K. (1982) Phil. Trans. Roy. Soc. Lond., A 305, p. 225-

247.

McKenzie D. (1978) Earth. Planet. Sci. Lett. , 40., £. 25-32.

Royden L. and Keen C.E. (1980) Earth Planet. Sci. Lett., 51^, _p_. 343-361.
N86-23129

Ill

TWO CONTRASTING METAMORPHOSED ULTRAMAFICMAFIC COMPLEXES FROM GREENSTONE BELTS, THE

NORTHERN KAAPVAAL CRATON AND THEIR SIGNIFICANCE IN ARCHAEAN TECTONICS

C.A. Smit and J.R. Vearncomfce, Dept. Geology, Rand Afrikaans

University, Johannesburg.

The character of Archaean ultramafic-mafic complexes can,

given their prominance in greenstone belts, provide critical

clues to help deduce the tectonic setting of these belts. Here

we describe two contrasting, metamorphosed, ultramafic-mafic

complexes, the first a partially serpentinised dunitic body with

associated chromite from Lemoenfontein, one of several peridotitic bodies occurring as discrete lenses and pods in granulite

facies gneisses of the northern Kaapvaal craton. The second,

the Rooiwater complex is a major layered igneous body, now metamorphosed in the amphibolite facies, but without pervasive deformation, which crops out in the northern Murchison greenstone

belt.

The Lemoenfontein body is circular, about 100m in diameter,

having the form of a steeply plunging boudin which complements

the regional structural pattern. The surrounding granulite facies gneisses were isotopically reset about 2650 Ma and may be

considerably older. The Lemoenfontein rocks are partially serpentinised dunite, displaying a prominent tectonic fabric defined by the preferred orientation of olivine grains, chromite

pods and disseminated chromite stringers, all of which are believed to have been through the granulite facies metamorphism.

Chromite is present as massive high-grade ore, 'leopard1 (nodular) ore, tectonically layered ore and disseminated ore. Zones

of chromite enrichment range in thickness from 1 to 30cm. The

Lemoenfontein chromites are similar to those mined in the Ultramafic Formation of the Selukwe greenstone belt, Zimbabwe.

Olivines from Lemoenfontein are Fo94 to Fo96 with NiO contents from 0.35 to 0.59wt%. The mineral chemistry of the chro-
' - CONTRASTING ULTRAMAFIC-MAFIC COMPLEXES

112 Smit, C.A. and Vearncombe, J.R.

mites of all different types (pods, trains, and inclusions in

silicate grains) is very similar indicating either complete metamorphic equilibration or they represent consistent primary

compositions. The Lemoenfontein chromites have refractory characteristics (low TiO2, A1203 and alkali metals) and plot on

geochemical fence diagrams in or close to the fields of other

podiform chromites. Rocks which in Phanerozoic series are closely associated with alpine-type peridotites or ophiolite suites.

The Rooiwater complex is a thick on end differentiated igneous body, of age greater than 2650 Ma, probably intruded at

2960 Ma. The complex is heterogenously deformed with much of

the 7.5km exposed thickness showing no pervasive deformation.

Metamorphosed pyroxenite, anorthosite, gabbro, sulphide-bearing

gabbros, thick magnetitite layers and differentiated granites

are compatible with the hypothesis that the body is a layered

intrusion although it is now allochthonous and intruded by

younger unrelated granites. Southward increasing TiO2 and decreasing V2O5 contents in magnetitite layers combined with a

general southerly disposition of differentiated hornblende granite suggest that the Rooiwater complex is southward facing. A

paucity of ultramafic cumulates and up to 1.5km of highly differentiated hornblende granite suggests that the original magma

was more felsic than that of similar layered intrusions.

The Lemoenfontein chromites and associated ultramafic rocks

are lithologically and chemically similar to their Phanerozoic

equivalents of ophiolitic origin, interpreted as obducted oceanic crust. Similarly we interpret the Lemoenfontein complex as

being a remnant of Archaean oceanic material. In contract, the

Rooiwater complex is, despite the lack of exposed intrusive contacts, similar to layered igneous complexes such as Ushushwana

or Bushveld. These complexes are intrusive in continental environments. We conclude that contrasting ultramafic-mafic complexes represent a heterogeneity in greenstone belts with either
CONTRASTING ULTRAMAF1C-MAFIC COMPLEXES

Smit, C.A. and Vearncombe, J,R,

oceanic or continental environments involved. Whether this

heterogeneity relates to a temporal or spatial (or both) control

remains uncertain.
114 N86-23130

ZIRCON Lu-Hf-SYSTEMATICS: '.EVIDENCE FOR "THE EPISODIC

DEVELOPMENT OF ARCHEAN GREENSTONE BELTS P. E. Smith, M. Tatsumoto* and R. M.

Farquhar; Dept. of Physics, Geophysics Division, University of Toronto. Canada

M5S 1A7; U.S. Geological Survey*, Federal Center, MS 963, P.O. Box 25046,

Lakewood, CO 80225

A combined U-Th-Pb and Lu-Hf isotopic study of zircons was undertaken

in order to determine the provenance and age of an Archean granite-greenstone

terrain and to test the detailed application of the Lu-llf system in various

Archean zircons.

The eastern Wawa subprovince of the Superior province consists of the

low grade Michipicoten and Gamitagama greenstone belts and the granitic

terrain. Earlier studies have established the structural and stratigraphic

relationships of the area (1-4). The adjacent high grade Kapuskasing zone

is believed to represent the lower crustal levels to the greenstone belts (5).

The rock units of this area have been the subject of extensive geochronological studies using zircon U-Pb (6, 7) and whole rock U-Th-Pb methods

(Smith, et al in prep.). The three volcanic cycles recognized in the area

have mean ages of 2748 My (cycle I), 2732 My (cycle II), and 2714 My (cycle

III). Syntectonic granitoids which surround the supracrustal rocks date

from the cessation of cycle I volcanic rocks, to the time of post-tectonic

plutonism dated at 2666 +_ 2 My. The oldest rocks yet dated come from a

granite dated at 2888 +_My which is possibly the basement to the volcanic

rocks. Zircon ages from the Kapuskasing zone appear to reflect updating

during the regional metamorphism (8).

The Lu, Hf, U and Th contents of zircons from these rocks reveal patterns

that may be indicative of their source regions (Fig. 1). Zircons from rocks

of granitic composition appear to have distinct enrichments in U and Th

relative to zircons from rocks of more intermediate composition. More

striking however, is the severe depletion of Lu and Hf from the zircons from

the Kapuskasing area. The lowest I If content measured so far, 1790 ppm, is

from zircons from a mafic gneiss. The elemental patterns in the lower

crustal zircons suggest that Lu and Hf loss accompanies Pb loss during high

grade metamorphism.

The U-Pb age corrected Hf isotopic ratios from the zircons indicate

significant long-lived heterogeneity of source regions for the greenstone

belts (Fig. 2). Overall the heterogeneity in the ratios may be attributed to

three isotopically distinct sources: (1) a high Lu/Hf source, (2) a moderately enriched Lu/Hf source; and (3) a sub-chondritic Lu/Hf source.

The high Lu/Hf source is represented by a sub-volcanic intrusive from

cycle II and two tholeiites (whole rock determinations) from the lower

stratigraphic levels of cycles I and II. The epsilon Hf values range from

+8.7 to +11.6 and the source is believed to represent the depleted mantle.

The second source has epsilon Hf values ranging from +1.4 to +5.9.

There is an apparent alignment of dacitic volcanic rocks and their subvolcanic equivalents from cycles I and II with the tonalitic syntectonic

granitoids. It is believed that the source of these rocks was the lower

crust and it can be inferred that previous intracrustal differentiation led

to a high Lu/Hf lower crustal reservoir. The process which led to the

enhanced Lu/Hf ratio was most likely Hf loss as attested to by the Kapuskasing zircons. A greater than chondritic Lu/Hf ratio for the lower crust may

explain the apparent non-coherence of initial Nd and initial Hf ratios for

an Archean tonalite reported in the literature (9, 10).
ZIRCON Lu-Hf SYSTEMATICS

Smith, P. E., Tatsumoto, M. and Farquhar, R. M. 115

The low Lu/Hf source is represented by rhyolites capping the sequences

of cycles I and III and by post-tectonic potassic granitoids. Their epsilon

Hf values ranging from -1.3 to +1.4, significantly lower than the coeval

dacites, are indicative of an upper crustal source.

The Hf isotopic data from the three volcanic cycles indicate that the

typical lithological features of a greenstone belt cycle could be acconmodated

in a crustal growth model that involved decreasing depth of melting in three

isotopically distinct reservoirs: mantle, lower crust and upper crust. The

model age of the sources given by the intersection of the lower crustal curve

with the bulk earth evolution curve (11) is about 2900 My, in good agreement

with the zircon U-Pb basement age. This linear array also has a similar intersection age to that of Proterozoic carbonatite complexes studied by Bell et

al (12) . The general convergence of the other reservoir vectors around this

age suggests that mantle depletion, crustal extraction and intracrustal

differentiation were all part of the same episodic event. It is also apparent

that recycling of older basement was important in the formation of many of

the later greenstone belt rocks.

Hf/100

Hf/100

Figure 1

Relative abundances of Lu, Hf, U and Th for eastern Wawa subprovince zircons.

Symbols are: A Dacitic volcanic rocks; v rhyolites; o sub-volcanic

granitoids; Q syntectonic granitoids; O post-tectonic granitoids; Q basement granite zircons; ft Kapuskasing zircons; • conglomerate boulder zircons.
116

ZIRCON Lu-Hf SYSTB1ATICS

Smith, P.E., Tatsumoto, M. and Farquhar, R. M.

02814 -

02809

2900 2850 2800 2750 2700

TIME (My)

2650 2600

Figure 2

Initial

 176Hf/177Hf vs T diagram for zircons and whole rocks ( D ). Symbols

as in Figure 1.

References

(1) Goodwin, A. M. (1962) Geol. Soc. Am. Bull. 73, 561-586.

(2) Ayres, L. D. (1969) Ont. Dept. Mines Geol. Rep. 69.

(3) Attoh, K. (1980) Current Res. part A. Geol. Surv. Can. Pap. 80-1 A,

101-106.

(4) Sage, R. P. (1980) Ont. Geol. Surv. Miscel. Pap. 96, 47-50.

(5) Percival, J.A. and Card, K. D. (1983) Geology 11, 323-326.

(6) Turek, A., Smith, P. E. and Van Schmus, W. R. (1982) Can. J. Earth Sci.

17, 866-876.

(7) Turek, A., Smith, P. E. and Van Schmus, W. R. (1984) Can. J. Earth Sci.

21, 457-464.

(8) Percival, J. A. and Krogh, T. E. (1983) Can. J. Earth Sci. 20, 830-843.

(9) Patchett, J. P., Kouvo, 0., Hedge, C. E. and Tatsumoto, M. (1981) Contr.

Miner. Petrol. 78, 279- 297.

(10) Jahn, B., Vidal, P. and Kroner, A. (1984) Contr. Miner. Petrol. 86,

398-408.

(11) Patchett, J. P. and Tatsumoto, M. (1980) Contr. Miner. Petrol. 75,

263-267.

(12)Bell, K. Blenkinsop, J., Cole, T. J. S. and Menagh, D. P. (1982) Nature

298, 251-253.
N86-S3131

EARLY PRECAMBRIAN CRUSTAL EVOLUTION OF SOUTH INDIA 117

*' '. - +•

R. SRINIVASAN, Geomysore Services, 12 Palace Road, Bangalore-560 052

The Early Precambrian sequence in Karnataka^ South India provides evidences

for a distinct trend of evolution which differs from trends exhibited in

many other Early Precambrian regions of the world. The supracrustal rock

associations preserved in greenstone belts and as inclusions in gneisses

and granulites suggest the evolution of the terrain from a stable to a mobile

regime. The stable regime is represented by 1. layered ultramafic-mafic

complexes, 2. orthoquartzite-basalt-rhyodacite-iron formation, and 3. orthoquartzite-carbonate-Mn-Fe formation. The mobile regime which can be shown on

sedimentol'ogical grounds to have succeeded the stable regime witnessed accumulation of a greywacke-pillow basalt-dacite-rhyolite-iron formation association.

Detrital sediments of the stable zone accumulated dominantly in fluvial environment and the associated volcanics are subaerial. The volcanics of the stable

regime are tholeiites derived from a zirconium and LREE-enriched source.

The greywackes of the mobile regime are turbidites, and the volcanic rocks

possess cbntinental margin (island-arc or back-arc) affinity; they show a

LREE-depleted to slightly LREE-enriched pattern. The evolution from a stable

to a mobile regime is in contrast to the trend seen in most other regions

of the world, where an early dominantly volcanic association of a mobile

regime gives way upward in the sequence to sediments characteristic of a

stable regime.

Structures in greenstone belts, in the gneisses surrounding them, and

also in the inclusions in the gneisses are similar in style, sequence, and

orientation. This structural unity which is present in spite of the three

thermal peaks recorded by radiometric ages around 3300, 3000 and 2600 m.y. ago,

indicates long range stability of tectonic stress regimes in the Archaean

lithosphere. The continuation of structures and rock formations across the

greenstone-granulite boundary suggests that the two provinces did not evolve

in separate tectonic blocks but 'represent only different crustal levels.

The preservation of detrital pyrite-uraninite bearing conglomerates,

iron formations, and carbonate rocks provide an unique opportunity for the

elucidation of evolutionary changes from oxygen-deficient to oxygenic atmospherehydrosphere conditions. Large scale development of iron formations and limestones in the greenstone belts of South India at least 3000 m.y. ago suggests

that these may be the earliest large-scale sinks for the photosynthetically

produced oxygen. Detailed palaeobiological and biogeochemical studies of

these rock formations are necessary.
118 N86-23132

BASEMENT-COVER RELATIONS AND INTERNAL STRUCTURE OF THE

CAPE SMITH KLIPPE: A 1.9 Ga GREENSTONE BELT IN NORTHERN QUEBEC,

CANADA; M.R. St-Onge and P.P. Hoffman, Geological Survey of Canada, 588

Booth St., Ottawa, Ont. K1A OE4, S.B. Lucas, Geological Sciences Dept, Brown

University, Providence, R.I. 02912, D.3. Scott, Geology Dept, McMaster University,

Hamilton, Ont. L8S 4L8, and N.3. Begin, Geological Sciences Dept, Queen's

University, Kingston, Ont. K7L 3N6.

The Cape Smith Belt is a 380x60 km tectonic klippe (1 and references therein)

composed of greenschist- to amphibolite-grade mafic and komatiitic lava flows and

fine-grained quartzose sediment, intruded by minor syn- to post-tectonic

granitoids. Previously studied transects in areas of relatively high structural level

show that the belt is constructed of seven or more north-dipping thrust sheets

which verge toward the Superior Province (Archean) foreland in the south and away

from an Archean basement massif (Kovik Antiform) external to the Trans-Hudson

Orogen (Early Proterozoic) in the north. A field project (mapping and structuralstratigraphic-metamorphic studies) directed by MRS was begun in 1985 aimed at

the structurally deeper levels of the belt and underlying basement, which are

superbly exposed in oblique cross-section (12 km minimum structural relief) at the

west-plunging eastern end of the belt. Mapping now complete of the eastern end of

the belt confirms that all of the metavolcanic and most of the metasedimentary

rocks are allochthonous with respect to the Archean basement, and that the thrusts

must have been rooted north of Kovik Antiform. The main findings (2) are:

1. A thin autochthonous to parautochthonous low-strain sedimentary sequence on

the south margin of the belt rests directly on Archean basement showing no

evidence of Proterozoic transposition.

2. The bulk of the belt is separated from the autochthon by a sole thrust which,

except at the south margin of the belt, is located at the basement-cover contact.

The hangingwall and footwall rocks of the sole thrust record high ductile strains

over a zone of increasing width, from south to north, toward the hinterland. Late

syn-metamorphic thrusts faults with relatively small displacements cut the sole

thrust and its associated shear zone, and place basement gneisses over cover rocks.

3. Lensoid meta-ultramafic tectonic blocks occur locally within the basal shear

zone. Their metamorphic anthophyllite-actinolite assemblage differs from the

serpentine-tremolite assemblage of cumulate meta-ultramafics occurring in sills at

higher structural levels. The blocks may have been tectonically transported from

mantle depths during thrusting, although this idea remains to be tested.

it. The allochthonous rocks above the sole thrust occur in a series of thrust sheets

bounded by south-verging (Dl) thrust faults, which are defined by structural

repetitions of stratigraphy and splay from the sole thrust. Favorable lithologies at

all structural levels (excepting the southern autochthonous margin) have a

pevetrative syn-metamorphic schistosity (SI) which is planar to south-facing tight

to isoclinal folds of bedding (Fl).

5. A transverse stretching lineation (LI) common in the lower stgructural levels

and pervasive in the basal shear zone, when considered with the Fl fold asymmetry

and overall thrust-ramp geometry, indicates relative southward translation of the

cover during Dl.
119

6. A pelitic interval above the sole thrust on the north margin of the belt contains

the metamorphic assemblage kyanite-staurolite-garnet-biotite-muscoviteplagioclase-quartz. The asemblage is indicative of metamorphic T of 550°C and

minimum P of 5.5 Kbars.

7. Mesoscopic late- to post-metamorphic chevron to rounded parallel folds (F2) of

the SI fabric have a marked limb asymmetry suggestive of a gravitational origin as

folds cascading off basement-cored macroscopic D2 antiforms into pinched coverrock synforms. The distribution of north- versus south-vergent mesoscopic folds

however is not always consistent with the mapped limbs of the macroscopic folds,

possibly reflecting diachronous development of the macroscopic folds.

8. Macroscopic high-angle D3 crossfolds affect both the basement and cover in the

eastern half of the belt and provide a cumulative structural relief of 12-15 km. D3

fold hinges are readily documented by reversals in plunge azimuth of the D2 folds.

Plunge projections permit the construction of a composite structural cross-section

linking the highest and lowest structural levels of the belt.

The main implication of these observations is that the presence of Archean

basement beneath the belt has no direct bearing on the question of the tectonic

setting of the mafic-ultramafic magmatism.

REFERENCES

(1) Hoffman, P.P. (1985) Can. 3. Earth Sci., 22, 1361-1369.

(2) St-Onge, M.R. et al. (1986), Geol. Surv. Can. Pap. 86-1A.
120

N86-23133

RHYOLITIC COMPONENTS OF THE MICHIPICOTEN GREENSTONE BELT, ONTARIO:

EVIDENCE FOR LATE ARCHEAN INTRACONTINENTAL RIFTS OR CONVERGENT PLATE MARGINS

IN THE CANADIAN SHIELD? Paul J. Sylvester, SM, NASA/Johnson Space Center,

Houston, TX 77058; Kodjo Attoh, Dept. of Geology, Hope College, Holland, MI

49423; Klaus J. Schulz, U.S. Geological Survey, Reston, VA 22092

Rhyolitic rocks often are the dominant felsic end member of the bimodal

volcanic suites that characterize many late Archean greenstone belts of the

Canadian Shield [1]. The rhyolites primarily are pyroclastic flows (ash flow

tuffs) emplaced following plinian eruptions [2], although deposits formed by

lava flows and phreatomagmatic eruptions also are present. Based both on

measured tectono-stratigraphic sections and provenance studies of greenstone

belt sedimentary sequences [3]» the rhyolites are believed to have been equal

in abundance to associated basaltic rocks.

In many recent discussions of the tectonic setting of late Archean

Canadian greenstone belts, rhyolites have been interpreted as products of

intracontinental rifting [2,M]. A study of the tectono-stratigraphic

relationships, rock associations and chemical characteristics of the

particularly well-exposed late Archean rhyolites of the Michipicoten

greenstone belt, Ontario (figure 1) suggests that convergent plate margin

models are more appropriate.

Three time-equivalent stratigraphic sequences of volcanism (figure 2),

each including both mafic and felsic rocks, have been recognized in the

Michipicoten greenstone belt [5,6,7»8]. The lower volcanic sequence is most

well-preserved and therefore has been studied in most detail. It consists of

a largely mafic unit (MV1) conformably overlain by a thick (up to about

700m), mainly felsic volcanic succession (FV1), which was emplaced

approximately 27^3 Ma ago [91. In the Michipicoten Harbour area, an undated

basal felsic flow unit is structurally discontinuous with the mafic sequence.

Along the northern margin of the belt, epiclastic sediments are deposited on

apparently older granitoid basement, and are overlain by felsic volcanics

(and iron formation) that may be time-correlative with the Michipicoten

Harbour felsic flows.

A range of depositional

environments apparently

existed for the felsic

volcanic rocks of the lower

volcanic sequence. Subaerial

non-welded massive ash flows,

shallow water accretionary

lapilli-bearing hyalotuffs

and deeper water bedded

pyroclastic deposits all have

been recognized [6,7»10].

Similarly, sedimentary rocks

that overlie the lower

volcanic sequence were

deposited in both subaerial

(braided fluvial and alluvial

fan) and subaqueous

(turbidite) environments

[11].

(Til >nccUIMt

Oorf form

TontfMc la QnnMc OiMto

Q3 FMng Olraclkm
MICHIPICOTEN RHYOLITES

Sylvester, P.J. et al.

ORIGINAL PAGE «

OF

 POOR QUALITY 121

107-

106-

110

111-

500

m

144

FV3

MV2

231B

\ FV1 172 .

1BT\ 422 . 197

 \ 419 .

416C -

• Voluminous Cenozoic rhyolitic

r:TM"" pyroclastlc deposits are erupted on

continental (rather than oceanic) crust

and exhibit distinctive chemical

characteristics and rock associations

depending on whether that crust was the

site of intracontlnental rifting or

subduction. Three examples of Cenozoic

rhyelites associated with

intracontinental, extension-related

tectonism are presented in Table 1.

The Trans-Pecos volcanic province of

west Texas represents a rift dominated

by alkaline to peralkaline rocks of

bimodal basalt-rhyelite composition.

The rhyolites are dominated by

low-silica (<75wt$) compositions that

tend to be depleted in alumina and lime

relative to iron and the alkalis. The

Rio Grande rift of New Mexico consists

of a more continuous spectrum of mafic

to felsic compositions that are

commonly described as calc-alkaline

[14]. Rhyolitic rocks, such as the

Bandelier Tuff, are dominated by high-silica compositions. The Yellowstone

Plateau volcanic field represents a third extension-related rhyolite group

characterized by an association with continental flood basalts and "hot spot"

277

£Z] MAFIC FLOWS

KT3 FELSIC

I

*-J PYROCLASTICS

R%| FELSIC FLOWS

^ IRON FORMATION

g?^] CONGLOMERATE

Fi^ OREYWACKE.

^""•^ SHALE, SILTSTONE

£23 VOLCANIC

^~~^ BRECCIA

tTT!) GRANITOID

343

Toimiions cuozoic IBIOLXTIC ije-FLOv TOFFS

Ixtaaalon-ralatad, Intraeoatiaantal «ult««

Trans Faooa Toloaole,

Frovlnoo, Tax** [123

taodallar Tuff, J«a»i

Hoaotalna, lav Naxloo

[13.1O

Tallovatoaa Flataau

Tolaaalo Flald [15]

Dominant

3i<>2 raoca

TO-T9 vt.l

T» - TT

T5 - TT

activity. Yellowstone rhyolites are compositionally

subalkaline rhyolites of the Rio Grande rift.

Cenozoic ash flow tuffs of

rhyolitic composition also are

erupted in voluminous

proportions in continental

inner arc regions of

convergent plate margins.

Relative to rhyolites formed

in intracontinental rifts or

hot spots, inner arc

subduction-related rhyolites

tend to have higher ratios of

alumina and lime to iron and

the alkalis (> about 1.4) and

a more continuous spectrum of

low- to high-silica

compositions. Three examples

of inner arc Cenozoic

rhyolites are listed in Table

2. They differ mainly with

respect to whether a field

association with voluminous

coeval intermediate volcanics

is present (San Juan field),

ambiguous (Sierra Madre

Occidental) or not found

similar to the

UljOj » CaO)/

(FsOt . alkali*)

O.T« - 1.S5

1.21 - t.lj

1.2« . 1.3T

Aaaooiatvtf

Toloaolea

batralit*,

•ugtarlt*.

tr.cbjt.

basaltic

aodaalt* to

rhyodaolta

oliTlna

thol.iit.

TOLOMIIOOS CIIOIOIC IHIOLITIC ISH'FLOV TOFFS

SnbdBotloo-ralatad, ooDtlaaetal lonar aro aaltaa

Taupo Toloanto Soaa,

lav laaland (1(1

Mld'Tartlar? Oppar

Toloaolo Saqaaaoa,

Slarra Madra Oooldaatal,

Maxloo (1T1

011(0can* lah-Flo«a,

San Joan Tolcanla Flald,

Colorado [It]

Doalnaot

3102 ranga

69-TT vt.J

TO - TT

6« - T6

(»1203 » CaO)/

(FaOt • alkalla)

1.5T - 1.T1

1.62 . 1.65

. 2.05

Aaaoelatad

Tolcaalca

alnor bl(h-ll

baaalt to

daolta

•laor baaaltlo

andaalta to

daolta

volualnoaa

andaalta to

qti latita
MICHIPICOTEN RHYOLITES

122 Sylvester, P.J. et al.

(Taupo volcanic zone). T»I.I s

If Cenozoic rhyolites may be LONU CICLI NZCBIMCOTII imoutij ir»n

used as a guide, the Michipicoten

lower volcanic sequence (FV1) sio'1?"*.

rhyolites, which are

, , j , , , . Lov-illloi tip* 70-73 ¥l.» 1.69 - *••• miner .nricb.d

characterized by a continuous D.I.U. <.dt.

spectrum of silica compositions m,h-.iiic. t,P. 7« - T« I.«T - z.2» -ioor .onob.<

and relatively high ratios of """Ul "c"*

alumina and lime to iron and the

alkalis (Table 3)» are more likely to be subduction-related than

intracontinental rift-related. The Taupo volcanic zone and neighboring

Kermadec-Tonga island arc system [19] offer perhaps the most appropriate

plate tectonic analogue. At this convergent plate margin, rhyolitic

pyroclastic rocks erupted from the New Zealand continental crust actually are

deposited largely on the adjacent sea floor [20], which also is the

depositional site for tholeiites derived from the Kermadec-Tonga island arc.

The resulting ocean floor/continental slope deposits should consist of

interfingering rhyolites and basalts derived independently from continental

and oceanic platforms, respectively.

A similar tectonlc-depositional model may explain the so-called cyclical

mafic to felsic stratigraphic relationships present in the Michipicoten belt.

The presence of pre-existing granitoid crust flanking the belt and the

well-known compositional similarity between Cenozoic island arc tholeiites

and Archean greenstone belt tholeiites [21], such as those present in the

Michipicoten belt [22], support this interpretation. However, the existence

of subaerial and shallow subaqueous depositional environments for some

Michipicoten volcanic, volcaniclastic and sedimentary units requires either

intermittent, local emergence of the volcanic pile or the existence of at

least small continental blocks underlying parts of the belt.

References: (1) Thurston et al. (1985) Geol. Assoc. Canada Spec. Pap.

28, p. 7-21. (2) Ayres, L.D. and Tburston, P.C. (1985) Ibid., p. 343-380.

(3) Ayres, L.D. (1983) Can. J. Earth Sci. 20, p. 1168-1194. (4) Gelinas, L.

and Ludden, J.N. (1984) Phys. Earth Planet. Int. 35. p.77-88. (5) Goodwin,

A.M. (1962) Geol. Soc. Am. Bull. 73, P- 561-586. (6) Attoh, K. (1980) Geol.

Surv. Canada, Paper 80-1 A, p. 101-106. (7) Attoh, K. (1981) Geol. Surv.

Canada, Paper 81-1B, p. 49-54. (8) Sage, R.P. (1979) Ontario Geol. Surv.,

Misc. Pap. 90, p.48-53. (9) Turek, A., Smith, P.E. and Van Schmus, W.R.

(1984) Can. J. Earth Sci. 21, p. 457-464. (10) Morton, H.L. and Nebel, M.

(1983) Precamb. Res. 20, p. 39-62. (11) OJakangas, R.W. (1983) LPI Tech.

Rpt. No. 83-03, p. 66-70. (12) Noble, D.C. and Parker, D.F. (1974) Bull.

Volcanol. 38, p. 803-827. (13) Sommer, M.A. and Schramm, L.S. (1983) Bull.

Volcanol. 46, p. 306-316. (14) Guilbeau, K.P. and Kudo, A.M. (1985) Geol.

Soc. Am. Bull. 96, p. 108-113. (15) Hildreth, W., Christiansen, R.L. and

O'Neil, J.R. (1984) J. Geophys. Res. 89, p. 8339-8369. (16) Cole, J.W. (1979)

N.Z. J. Geol. Geophys. 22, p. 631-657* (17) Cameron, K.L. and Hanson, G.N.

(1982) Geochim. Cosmochim. Acta 46, p. 1489-1503. (18) Lipman, P.W. (1975)

U.S. Geol. Surv. Prof. Pap. 852, 128pp. (19) Ewart, A., Brothers, R.N. and

Mateen, A. (1977) J. Volcanol. Geotherm. Res. 2, p. 205-250. (20) Walker,

G.P.L. (1980) Ibid., 8, p. 69-94. (21) Condie, K.C. (1985) J. Petrol. 26, p.

545-563. (22) Sylvester. P., Attoh, K. and Schulz, K.J. (1985) Geol. Assoc.

Canada Program with Abstr. 10, p. A60.
N86-23134 123

GEOPHYSICAL CHARACTERISTICS AND CRUSTAL STRUCTURE OF GREENSTONE

TERRANES, CANADIAN SHIELD; M.D. Thomas1, L. Losier2, P.C.

Thurston3, V.K. Gupta3, R.A. Gibb1 and R.A.F. Grieve1,

1Earth Physics Branch, EMR, Ottawa; 2Geological Sciences, McGill

University, Montreal;

 3Ontario Geological Survey, Toronto.

A knowledge of the deep structure and geometry of greenstone belts is

fundamental to tectonic models of Archean evolution. In the Canadian

Shield long linear granite-greenstone terranes of generally low

metamorphic grade alternate with temporally-equivalent metasedimentary

belts of higher grade. The focus of geophysical investigations of these

terranes has been to examine geometries and contact relationships within

individual terranes, and to look at the broader and deeper aspects of

structure and inter-terrane relationships.

Major greenstone belts are characterized by positive gravity

anomalies in the range 15-30 mGal that primarily reflect the relatively

high density mafic and ultramafic metavolcanic components (1). These

anomalies are sometimes interrupted by negative anomalies caused by

felsic plutons and are poorly developed where high metamorphic grade

basement is present and/or boundaries are gently-dipping. Modelling

reveals that many greenstone belts are more or less basin-shaped, some

having deep keels, and that their steep surface boundaries extend to

depth. Model depths of polycyclic greenstones are 2-8 km and

non-polycyclic are 3-12 km (1). The generally smaller depths of the

former have been attributed to granitic intrusion decreasing vertical

extent by stoping, or to listric normal faulting or thrusting (1).

Models indicate abrupt changes in depth of up to ~10 km between

supracrustals of the Wawa greenstone and Quetico metasedimentary terranes

and point to a major faulted contact (2). Granitic intrusions at and

within boundaries of greenstones are associated with prominent negative

gravity anomalies. Modelling indicates that they have depths ranging

from 2-16 km with depths in the middle of the range being characteristic

(3,4). Generally, the contacts of the granites are modelled as steeply

dipping. Some granites extend several kilometres deeper than adjacent

greenstones but in other cases greenstones are interpreted to underlie

the granite. For example, interpretation of a combined gravity-seismic

study of the Aulneau batholith of the Wabigoon subprovince suggests that

it is floored by up to 10 km of greenstones (3). Gravity studies in

Wabigoon subprovince have contributed to classifying granites into

epizonal sheets and deep diapiric batholiths intruded in two separate

periods (4).

Regionally, greenstone belts generally correspond to magnetic lows

and associated granites to magnetic highs (5,6). Magnetization studies

(6) indicate values that are generally < 0.05 A/m for greenstones and

> 0.05 A/m for granites. Linear positive anomalies within the English

River gneiss belt have drawn attention to pyroxene amphibolite gneisses,

probably derived from metavolcanics (7). Their occurrence is significant

in that they are in an area where volcanism is thought not to have been

important. Aeromagnetic shaded relief maps have been used to assist in

mapping surface geology in the Abitibi greenstone belt (8). Various

features correlate with diorite-gabbro and peridotite-serpentinite

intrusions, diabase dykes, major faults, iron formations and zones of

contact metamorphism around granitic intrusions. The magnetic signature
raA

124 * Thomas, M.D. et al.

of the Abitibi belt, however, is not noticeably different from that of

the bordering terranes. Modelling has been limited. Interpretation of a

300 km N-S profile across the Abitibi belt (8) indicates that the

greenstones extend to a maximum depth of 13.6 km in the south, with an

average depth of ~9 km compared to 6 km in the north. This agrees with

seismic refraction results that suggest the bottom of the belt dips

southward increasing in depth from 6 to 14 km (9). Surface magnetic

units over granites of the Wabigoon belt have been modelled as extending

to the intermediate discontinuity (16-19 km) with an increase in

magnetization occurring at a few kilometres depth (6). Magnetization is

low or absent below the discontinuity.

Seismic reflection studies within the Aulneau batholith and adjacent

greenstones (10,11) have mapped a near-vertical contact between granite

and greenstone to a depth of several kilometres (confirmed by later

gravity studies) and a vertical fault zone. Although there are no

detectable velocity differences between the greenstones and granites, the

impedance contrast is sufficient to produce recognizable reflections from

the near-vertical contact. The lower surface of the batholith, as

interpreted from gravity, did not produce reflections, perhaps due to

its undulatory nature (12). There is also a poor correlation between the

average depth of the Yellowknife greenstone belt as determined from

seismic (~10 km) and gravity (~3 km) studies (13,14). In contrast,

the seismic refraction survey (9) across the Abitibi belt yielded a

geometry for the bottom of the belt similar to that based on magnetic

interpretation (8). The seismic investigations in the vicinity of the

Aulneau batholith (10,11) also detected several deep horizontal or

near-horizontal reflectors. The most prominent reflectors are at

intermediate depths of about 19 and 22 km and the Moho at 38 km. The

three reflectors appear to be continuous beneath the granite and

greenstones suggesting that complex structure, which typifies the upper

crust, is absent at depth. A similar picture of the Wabigoon crust has

been found by long-range refraction - wide angle reflection experiments

(15,16), but in the Quetico metasedimentary belt to the south no sharp

boundaries are found within or at the base of the crust which is about

40-42 km thick (16). In the English River gneiss belt to the north

seismic refraction studies indicate thinner crust with an average

thickness of 34 km (17). The average depth of the intermediate

discontinuity remains about the same (~18 km). In detail, the Moho is

upwarped by roughly 8 km in the northern part of the belt, whereas the

intermediate discontinuity exhibits a complementary downwarp with an

amplitude of 10 km. Re-examination of the orginal data (12) indicates

that the axis of this proposed warping lies close to the northern margin

of the gneiss belt where it coincides with a sedimentary basin.

Magnetotellurie investigations have been carried out in the western

Wabigoon belt (18). A 3.9 km thick near-surface resistive zone under the

metavolcanics is considerably less resistive (21,300 ft-m) than one 7.4

km thick under the granitic gneiss (3,280,000 fl-m). It suggests that

crust underlying metavolcanic rocks is partially fractured and contains

saline fluids and/or that the metavolcanics extend throughout the

resistive zone. Heat flow studies reported from several Precambrian

shields indicate that the average heat flow in greenstones is roughly 10%

lower than in crystalline terranes (19). Heat generation data from the

Churchill and Superior provinces of the Canadian Shield indicate that
GEOPHYSICAL CHARACTERISTICS

Thomas, M.D. et al.

 125

greenstones are approximately 7 km thick.

A general conclusion is that greenstone belts are not rooted in deep

crustal structures. Geophysical techniques consistently indicate that

greenstones are restricted to the uppermost 10 km or so of crust and are

underlain by geophysically normal crust. Gravity models suggest that

granitic elements are similarly restricted, although magnetic modelling

suggests possible downward extension to the intermediate discontinuity

around ~18 km. Seismic evidence demonstrates that steeply-dipping

structure, which can be associated with the belts in the upper crust, is

not present in the lower crust. Horizontal intermediate discontinuities

mapped under adjacent greenstone and granitic components are not

noticeably disrupted in the boundary zone. Geophysical evidence points

to the presence of discontinuities between greenstone-granite and

adjacent metasedimentary terranes. Measured stratigraphic thicknesses of

greenstone belts are often twice or more the vertical thicknesses

determined from gravity modelling. Explanations advanced for the

discrepancy include stratigraphy repeated by thrust faulting and/or

listric normal faulting (1), mechanisms which are consistent with certain

aspects of conceptual models of greenstone development. Where repetition

is not a factor the gravity evidence points to removal of the root zones

of greenstone belts. For one region, this has been attributed to

magmatic stoping during resurgent caldera activity (20).

Geophysical studies in the Canadian Shield have provided some

insights into the tectonic setting of greenstone belts. Much work,

however, remains to be done, particularly in the use of geophysics in

evolutionary models of greenstone development. Future needs include

detailed, integrated studies, the introduction of relatively new methods

such as Vibroseis seismic reflection, greater use of magnetotellurics and

the application of other electromagnetic methods such as very low

frequency (VLF) surveys.

References

(1) Gupta V. K. and Thurston P.O. 1985. Geology (in preparation).. (2)

Cheadle S.P. 1982. MSc Thesis, Lakehead University. (3) Brisbin W.C.

and Green A.G. 1980. C.J.E.S., 17, 968-977. (4) Szewczyk Z.J. and West

G.F. 1976. C.J.E.S., 13, 1119-1130. (5) Magnetic Anomaly Map of Canada.

1984. G.S.C. Hap 1255A. (6) Hall D.H. 1968. C.J.E.S., 5, 1277-1296.

(7) Urquhart W.E.S. 1977. Geotraverse Conference Abstract, 5-7. (8)

Losier L. 1984. MSc Thesis, Ecole Polytechnique de Montreal. (9) Parker

C.L. 1984. MSc Thesis, McGill University. (10) Green A.G., et al.,

1978. C.J.E.S., 15, 301-315. (11) Green A.G., et al., 1979. C.J.E.S.,

16, 1599-1612. (12) Hall D.H. and Brisbin W.C. 1982. C.J.E.S., 19,

2049-2059. (13) Glee, T.E., et al. , 1974. C.J.E.S., 11, 1534-1549.

(14) Gibb R.A. and Thomas M.D. 1980. C.J.E.S., 17, 1506-1516. (15)

Hall D.H. and Hajnal Z. 1973. Bull. Seis. Soc. Am., 63, 885-910. (16)

Wright J. and West G.F. 1976. Geotraverse Conference, 47-55. (17) Hall

D.H. and Hajnal Z. 1969. C.J.E.S., 6, 81-99. (18) Kryzan A. and

Strangway D.W. 1977. Geotraverse Conference, 121-131. (19) Drury M.

1985. Abstract (this volume). (20) Gupta V.K. et al., 1982. Precam.

Res., 16, 233-255.
126 N86-23135

VOLCANOLOGICAL CONSTRAINTS ON ARCHEAN TECTONICS P. C. Thurston and

L.D. Ayres, Ontario Geological Survey, Toronto, Ont., Canada, M5S 1B3; Dept.of

Earth Science, Univ.of Manitoba, Winnipeg, Man., Canada. R3T 2N2.

Terranes within Archean shields can be classified as granite-greenstone

megabelts, contemporaneous sedimentary megabelts (1) and basement enclaves

within either of the above (2,3). Stratigraphic and geochronological work in

the Superior Province has shown the granite-greenstone megabelts represent

proximal volcanism, sparse deep water clastic sedimentation, and late alluvial

fan-submarine fan sedimentation(4). The sedimentary megabelts represent

stratigraphically equivalent deep water sedimentation of wacke-pelite couplets

(5), submarine fan conglomerates and minor distal facies volcanismt6). Basement

enclaves include meta-igneous and metasedimentary gneiss and fragmented

metavolcanic relics with poorly preserved primary textures.(2,7).

Greenstone belts of Australia have been subdivided into >3 Ga platform-phase

greenstones and <3Ga rift-phase greenstones (8). Platform phase units are basal

komatiite flows and tholeiitic flows with an upper unit of minor pyroclastics.

Volcanism in the platform phase is typified by abundant pillowed amygdular

flows, overlain by minor airfall tuff and relatively distal debris flow

volcaniclastic units. Sedimentary units include chert, quartzite, and

stromatolitic carbonates with minor wackes, indicative of shallow water

platform sedimentation (9). Examples in the Superior Province, generally about

3 Ga old (10) include quartz-rich wackes in the lower sequence at North Spirit

Lake in the Sachigo Subprovince (11), quartzites with fuchsite clasts in the

lower sequence of the Wabigoon Subprovince at Armit Lake (12), and carbonaterich sediments in the Lumby Lake greenstone belt (13) within the Wabigoon

Diapiric Axis basement enclave (3). Volcanologically one can conclude from the

thickness of the shallow water volcanic rocks and sediments that accumulation

took place on a shallow platform (9) and as well, large scale subsidence kept

pace with the rate of accumulation of volcanic rocks.

Rift-phase greenstones are relatively deep water amygdule-poorlpillowed

tholeiites succeeded upward by vesiculated pillowed flows and calc-alkaline

pyroclastic and volcanoclastic units (8). Considering the maximum water depth

for pyroclastic erutions (14) and the thickness of pyroclastic sections in many

rift-phase greenstone belts, Ayres (15) has suggested many Plinian eruption

columns became subaerial. Classically (15) most Archean pyroclastic units were

considered to have been deposited subaqueously. Recent studies have shown

however that many Archean pyroclastic units were deposited subaerially (16,15).

Sedimentologic studies of rift-phase greenstones show some deep-water clastic

deposits(17), but increasingly shallower water deposits (alluvial fan) at

stratigraphically high levels.' The structural pattern in rift-phase belts is

alternating synclinoria and anticlinoria either breached by diapirism or

sheared out (8). Most Superior Province greenstone belts younger than 2.9-2.8

Ga (18) are probably rift-phase based on the following, a)structural style with

synclinoria dominating with only rare dome and basin structures, 'b)

Quartz-rich and carbonate rich sedimentation is scarce in the Abitibi (4),

kWabigoon(20), and younger (<2.9Ga) sequences of the Uchi (18) and Sachigo

(11,18) Subprovinces. c) Volcanism is typified by bimodal tholeiite-rhyolite

sequences (21) with shoaling upward attributes (18,4). Evidence for small

scale operation of rift-related volcanism is seen in the Six Mile Lake cycle at

Sturgeon Lake (22) where a tholeiitic basalt-calcalkaline rhyolite sequence

2755 Ma (23) is rich in incompatible elements relative to later sequences and

is cut by abundant mafic dikes assumed to feed the uounger (2718Ma) (24) cycles

related to wide-scale rifting.

A survey of volcanic cyclicity (20) reveals the following types of cycles

within the Superior Province. (+=fractionation relation; -=no fractionation

KOM =komatiitic; TH=tholeiitic; CA=calc-alkaline; ALK=alkaline magma clans

1)KOM Perid Kom +dacite 4)TH bas -I- andes - Ca bas + rhy - ALK

2)KOM Perid.kom - TH bas + rhy - 5)CA bas + rhy

-CA bas + rhy -ALK

3)TH bas + andes - Th andes - Ca dac 6)TH bas - CA dac + rhy - THbas

+rhy

Increasing Stratigraphic height to the right in each entry.

Cycle types 3,4, and 5 above were formerly thought to represent

fractionation sequences/but recent work has shown that many are bimodal(21).

The fact that the above cycle types are bimodal has profound volcanologic and

petrogenetic implications in that the bimodalism is not simply the paucity of

intermediate composition magmatic liquids. Trace element geochemistry and

field evidence suggests, when corrected for unerupted volume in zoned magma
VOLCANOLOGIC CONSTRAINTS 127

Thurston,P.C. and Ayres, L.D.

chambers,and loss of vitric fines in high level winds during Plinian eruptions

are made, preserved volumes of felsic volcanics in the Archean represent +15%

of the original felsic magma (21). In effect, we concluded that Archean

bimodal volcanism represents subequal volumes of mafic and felsic magma which

are involved in greenstone belt volcanism.

Determination of paleoenvironment (above), eruption type, eruption rate,

magma chamber size and type, developmental processes, and the life span of

individual volcanoes places many genetic constraints on greenstone belt

tectonics. In mafic sequences subequal volumes of pillowed and massive flows

(18) suggest eruption by sheet flow processes (25) dominate over eruption from

shield volcanoes (18). In felsic sequences the volumetric dominance of

ignimbrites (21) and the notion that sedimentary basins contain large amounts

of tephra suggest Plinian eruptions were dominant in the Archean. Many Plinian

eruptions produced subaerial deposits on local volcanic islands (18.19,15).

Vulcanian eruptions are subordinate, they produce less widespread deposits -

examples include the Skead Group (26) and the Lake of the Woods area (27).

This eruption type is often the result of less volatile-rich magmas relative to

Plinian systems (28) interacting with near-surface water. The deposits are

generally less widespread in extent than many Plinian deposits.

Eruption rates of Archean volcanoes can be determined in an approximate

and indirect fashion. Sheet flows (25) a greater mean flow thickness than in

Phanerozoic analogues (18) and the presence of lava plains (29) in Archean

mafic sequences suggest a more rapid eruption rate than in Phanerozoic

analogues (30). Phanerozoic ignimbrite systems have volumes in the 10^-10^ km3

range (31) with exceptional examples in the 103-104 km3 range (31,32).

Phanerozoic felsic volcanoes had a life-span generally not exceeding 1.5 Ma

(18) but many Archean felsic edifices apparently existed for 10-20 Ma(18).

The preserved volume of felsic ignimbrites (recalculated to compensate for

unerupted material and loss of vitric fines, but ignoring compaction)

suggests existance of felsic magma chambers on the order of 103 km3 (21)

rivalling those of the largest Phanerozoic systems (28,29). When integrated

with data on the lifespan of Archean volcanoes of 10-20 Ma, Archean felsic

eruption rates were large, but not as large as those seen in Archean mafic

systems.

Volcanological and trace element geochemical data can be integrated to place '

some constraints upon the size, character and evolutionary history of Archean

volcanic plumbing, and hence indirectly, Archean tectonics. The earliest

volcanism in any greenstone belt is almost universally tholeitic basalt.

Archean mafic magma chambers were usually the site of low pressure

fractionation of olivine, plagioclase and later Cpx+_ an oxide phase during

evolution of tholeitic liquids (33 and references therein). Several models

suggest basalt becoming more contaminated by sial with time (33,34). Data in

the Uchi Subprovince shows early felsic volcanics to have fractionated REE

patterns (33)followed by flat REE pattern rhyolites. This is interpreted as

initial felsic liquids produced by melting of a garnetiferous mafic source

followed by large scale melting of LIL-rich sial (33). Rare andesites in the

Uchi Subprovince are produced by basalt fractionation, direct mantle melts and

mixing of basaltic and tonalitic liquids(33). Composite dikes in the Abitibi

Subprovince (35) have a basaltic edge with a chill margin, a rhyolitic interior

with no basalt-rhyolite chill margin and partially melted sialic inclusions.

Ignimbrites in the Uchi (16) and Abitibi (36) Subprovinces have mafic pumice

toward the top. Integration of these data suggest initial mantle-derived

basaltic liquids pond in a sialic crust, fractionate and melt sial. The

inirial melts low in heavy REE are melts of mafic material, subsequently

melting of adjacent sial produces a chamber with a felsic upper part underlain

by mafic magma.

Compositional zonation of the overlying felsic magma develops with time (31),

resulting in Plinian eruption through roll over (37) or volatile

supersaturation(38).

Numerous arguments suggest widespread volcanism-related subsidence kept pace

with the rate of eruption: a) The preservation of felsic sequences rather than

the rapid erosion common in Phanerozoic terranes (39) b)Minimum water depth for

pyroclastic activity (14) vs preserved stratigraphic thickness of subaqueous

pyroclastic units (15)i.e. sections are much thicker than maximum water depth
128 VOLCANOLOGICAL CONSTRAINTS

Thurston P.C., and Ayres L.D.

for eruption - therefore subsidence occurred. c) Lateral extent of 30-50 km

for stromatolitic carbonates (40) in the Uchi subprovince,lateral extent of

30-50 km for shallow water silicified evaporites (41) and lateral extent and

high eruption rate for shallow water environment mafic plains would have

rapidly become subaerial unless subsidence kept pace(18). Isostatic

calculations (42,43) suggest lava plain eruptions produce lesser crustal

loading than central vent eruptions and less isostatic subsidence. Models

involving sialic substrate to lava plain systems produce (42) sufficient

subsidence to just maintain volcanic piles at sea level. Therefore we conclude

a) subsidence kept pace with volcanism, b)subsidence was regional in extent, c)

it is difficult to envision a sagduction style of subsidence (44) producing

subsidence over a large area consistent with the great areal extent of the main

contributor to the subsidence- the mafic lava plains. Subsidence was more rapid

during mafic volcanism slowing during felsic volcanism.

The great volumes of Archean rhyolites and bimodal nature of rift-phase

volcanism mitigates against an island arc or back-arc basin analogue where

rhyolite is scarce (39 and references therin). Both continental arcs and

continental rifts have sufficient volumes of felsic volcanism to compare to

greenstone belts. The sediment-filled grabens associated with the Rio Grande

Rift (45) offer a possible modern analogue as do the continental intra-arc

depressions (39).

REFERENCES

DYoung GM 1978, Geosci. Can. 5:140-149.

2)Beakhouse GP 1985, Geol Assoc Can

Spec pap 28:169-178

3)Thurston PC and Davis DW 1985,

OGS MP 126

4)Dimroth et al 1982, Can J Earth Sci

19:1729-1758

5)Thurston PC 1985,OGS GR 234

6)Bowen RP 1984 OGS OFR 5471 '

7)Thurston PC and Davis DW OGS

MP 126

8)Groves DI and Batt WD 1984 Archaean

Geochemistry, Springer 73-98

9)Eriksson KA 1981 Geol Soc Aust Spec

Pub 7:235-244

10)Corfu et al 1984 Geol Ass'n Can

Abst 9:54

IDWood J 1980 Prec Res 12:227-255

12)Sutcliffe RH 1985 pers comm

13)Jackson MC 1984 OGS OFR 5534

14)McBirney AR 1963 Bull Vole

26:455-469

15)Ayres LD 1982 Geol Ass'n Can

Short Crs Notes 2:1-17

16)Thurston PC Prec Res 12:79-98

17)0jakangas RW 1985 Geol Ass'n Can

Spec Pap 28:23-48

18)Ayres LD and Thurston PC 1985 Geol

Ass'n Can Spec Pap 28:343-380

19)Dimroth et al 1983 Can J Earth Sci

20:1374-1388

20)Blackburn et al 1985 GeolAss'n Can

Spec Pap 28:89-116

21) Thurston et al 1985 Geol Ass'n Can

Spec Pap 28:7-21

22)Trowell NF 1983 OGS GR 221

23)Davis DW 1985 pers comm

24)Davis DW and Trowell NF 1982 Can J

Earth Sci 19:868-877

25)Greeley R 1982 J Geophys Res 87:2705

-2712

26)Jensen LS 1985 Geol Ass'n Can Spec

Pap 28:65-88

27)Car D 1976 Ctr Prec Stud U of Man.

89-96

28)Fisher RV and Schminke H-U 1984

Pyroclastic Rocks, Springer

29)Dimroth E et al 1984 Geol Ass'n Can

Spec Pap 28:203-222

30)Wadge G 1982 J Geophys Res 87:4035-

4049

31)Smith RL 1979 Geol Soc Amer Spec Pap

180:5-27

32)Lipman PW 1984: J Geophys Res 89:

8801-8841

33)Thurston PC and Fryer BJ 1983 Contr

Min Pet 83:204-226

34)Gelinas L and Ludden JN 1984 Phys

Earth & Planet Int 35:77-88

35)Gibson H 1983 pers comm

36)Thurston PC 1985 OGS MP 126

37)Rice A 1981 J Geophys Res 86:405-417

38)Sparks RSJ et al 1977 Nature 267:315

-318

39)Hildebrand RS and Bowring SA 1984

Geology 12:73-77

40)Hofmann et al 1985 Geol Ass'n Can

Spec Pap 28:125-132

41)Fisher GL and Lowe DR GSA Abst 17:1:

583

42)Dimroth E et al 1984 CIMM Spec Vol

34:73-91

43)van Wagoner N et al 1985 McGill U

Absts 7

44)Goodwin AM 1982 Can J Earth Sci

19:1140-1155

45)Chapin CE 1979 Rio Grande Rift:

Tectonics & Magmatism 1-5
N86-23136

129

THE DEHYDRATION, REHYDRATION AND TECTONIC SETTING OF GREENSTONE BELTS

IN A PORTION OF THE NORTHERN KAAPVAAL CRATON, SOUTH AFRICA; D.D. van

Reenen, J.M. Barton Jr., C. Roering, J.C. van Schalkwyk, C.A. Smit

(Dept. of Geology, RAU, P.O. Box 524, Johannesburg 2000, S. Africa)

J.H. de Beer (NPRL, CSIR, P.O. Box 395, Pretoria 0001, S. Africa) and

E.H. Stettler (Geological Survey, Private Bag X112, Pretoria 0001, S.

Africa).

High-grade gneiss terranes and low-grade granite-greenstone terranes are

well known in several Archaean domains. The geological relationship between

these different crustal regions, however, is still controversial. One school

of thought favors fundamental genetic differences between high-grade and lowgrade terranes while others argue for a depth-controlled crustal evolution.

The detailed examination of well-exposed Archaean terranes at different metamorphic grades, therefore, is not only an important source of information

about the crustal levels exposed, but also is critical to the understanding

of the possible tectonic and metamorphic evolution of greenstone belts with

time.

Many features of a metamorphic and deformational transition from a typical greenschist facies granite-greenstone terrane to a high-grade gneiss terrane are illustrated in the crustal section of the northern portion of the

Kaapvaal Craton over the 60 km between the Pietersburg greenstone belt and

the granulite facies Southern Marginal Zone of the Limpopo Belt. In this

section, steeply-dipping, typical greenstone belt lithologies occur at higher

and higher grade moving from south to north. In the south, the Pietersburg

belt comprises an at least 3450 Ma mafic, felsic and ultramafic volcanic and

volcano-sedimentary assemblage (the Pietersburg Group) unconformably overlain

by a sedimentary assemblage (the Uitkyk Formation), probably deposited between about 2800 Ma and 2650 Ma. The Pietersburg Group is surrounded by the

approximately 3500 Ma tonalitic and trondhjemitic Baviaanskloof Gneiss and is

intruded by the approximately 2800 Ma Hout River Gneiss. The belt is apparently intruded by approximately 2650 Ma, largely undeformed, granodioritic

plutons. Metamorphic grade within the Pietersburg belt increases from greenschist facies in the southwestern and central parts to amphibolite facies in

the northeast, consistent with the regional metamorphic pattern.

North of the Pietersburg belt are situated mafic, felsic and ultramafic

volcanic and sedimentary rocks of the Rhenosterkoppies greenstone belt. The

Sutherland greenstone belt, composed of mafic and ultramafic and mafic-tofelsic volcanic rocks and sedimentary rocks, occurs to the east. Both belts

are surrounded by the Baviaanskloof Gneiss. The ages of the lithologies

within the Rhenosterkoppies and Sutherland belts are unknown, but both have

been metamorphosed under amphibolite facies conditions.

In the Southern Marginal Zone, highly attenuated and boudinaged granulite facies greenstone belt lithologies (mafic, ultramafic and metapelitic

gneisses and banded iron formation) occur within the Baviaanskloof Gneiss.

These assemblages are intruded by the approximately 2650 Ma deformed Matok

charnockitic-granodioritic pluton while the undeformed Palmietfontein granite

was emplaced at about 2450 Ma.

The transition from the lower-grade granite-greenstone terrane to the

Southern Marginal Zone is not only refelcted by an increase in the grade of

metamorphism but also by an increase in the intensity of deformation. The

structural grain of the entire area trends roughly east-northeast with an

almost vertically dipping schistosity or gneissosity. In the Southern Mar-
DEHYDRATION, REHYDRATION AND TECTONIC SETTING

130 D.D. van Reenen et al.

ginal Zone, the distended nature of the granulitic greenstone remnants is in

sharp contrast to the more continuous outcrop pattern of the greenstone lithologies to the south. The magnetic lineation patterns of the area change

abruptly at the boundary of the granulite facies terrane and seismic velocities increase sharply at the same point, possibly as a result of mantle material being at a higher level.

The crustal behavior of the entire region must have been consistent with

the observation that the rocks of the Southern Marginal Zone were depressed

into deep crustal levels. This movement probably implies that the lowergrade terranes were depressed in a sympathetic manner. In the Southern Marginal Zone, the maximum prograde conditions (P>9.5 kb and T>800°C) reached

during this tectonic event are recorded by the assemblage garnet + hypersthene

+ quartz + plagioclase +/- biotite +/- kyanite in metapelite. These conditions were followed by rapid, nearly isothermal, decompression between approximately 2700 Ma and 2650 Ma, recorded by decompression textures of cordierite and hypersthene after garnet. P-T conditions of this decompression

event were T=800°C with P decreasing to 7.0 kb. The Matok pluton was emplaced toward the end of the isothermal decompression. The southern margin

of this dehydrated terrane was then subjected to a regional encroachement of

C0_-rich hydrating fluids before approximately 2450 Ma, the time of emplacement of the Palmietfontein granite. This infiltration produced the retrograde orthoamphibole isograd defined by the reactions: hypersthene + quartz

+ H20 = anthophyllite and cordierite + H20 = gedrite + kyanite + quartz.

These reactions occurred at T=650°C to 600°C and a total P less than 6 kb at

PH-0 = 0.2 P ,. Completely hydrated and recrystallized rocks south of

this isograd are characterized by the assemblage anthophyllite + gedrite +

kyanite + biotite + quartz + plagioclase. The fluids responsible for rehydration are believed to have been derived from hydrated granite-greenstone

lithologies.

Metamorphic assemblages in the area south of the retrograde isograd are

still insufficiently documented to delineate isograds but the overall increase

in the grade of metamorphism from south to north is illustrated by comparing

assemblages in chemically similar granite-greenstone lithologies from different crustal levels:

LITHOLOGIES PIETERSBURG BELT RHENOSTERXOPPIES SOUTHERN MARGINAL

(Central/South AND SUTHERLAND ZONE OF LIMPOPO

West part) BELTS BELT

Meta-sediments Qz + Chi • + f eldsp + Orthoamph + Blot + Hyp + Blot + Qz +

sericite ± Carb Qz + Plag + Garn Flag «• Garn + cord

* Kyan

Mafic Rocks Ab + Ep + Chi + Act Hbl + Plag (An 33) Opx + Cpx * Plag

+ Qz + Carb + Qz ± Sf ± Diop (An 49) + Mt ± Hbl

i Qz

Ultramafic Rocks Trem + Chi i Talc Trem + Chi i Ol Ol » Opx + Sp + Call Carb ± Carb Amph
DEHYDRATION, REHYDRATION AND TECTONIC SETTING

D.D. van Reenen et al.

Gravity and resistivity data indicate that the Pietersburg, Rhenosterkoppies and Sutherland greenstone belts are shallow crustal features, rarely

exceeding 5 km in depth. This depth is in marked contrast to the thickness

of the various lithologic successions measured across the stratification.

These observations plus the fact that the crustal thicknesses determined from

stratigraphic sections are far in excess of those which can be accepted from

the metamorphic data, suggest that major crustal thickening took place in

this area by tectonic stacking. A steeply northward dipping shear zone with

near vertically plunging mineral lineations is exposed along the Hout River.

This shear zone is believed to be associated with the tectonic stacking but,

as yet, the sole thrust for this proposed stacking has not been recognized.

The hydrating fluids responsible for establishing the retrograde orthoamphibole isograd could have been derived by dehydration of over-ridden lower-grade

crustal rocks during and after thrusting.
132 N86-23137

THE STRATIGRAPHY OF THE STEEP ROCK GROUP, N.W. ONTARIO, WITH EVIDENCE OF A

MAJOR UNCONFORMITY

M.E. Wilks and E.G. Nisbet

Dept. of Geological Sciences, University of Saskatchewan, Saskatoon

The Steep Rock Group is exposed 6 km north of Atikokan, 200 km west of

Thunder Bay. It is situated on the southern margin of the Wabigoon Belt of

the Archaean Superior Province, N. W. Ontario. Reinvestigation of the

geology of the Group has shown that the Group lies unconformably on the

Tonalite Complex to the east.

This unconformity.has been previously suspected, from regional and mine

mapping but no conclusive outcrop evidence for its existence has as yet been

published.

The strike of the Group, comprised of Basal Conglomerate, Carbonate Member,

Ore Zone and Ashrock is generally north-northwest dipping steeply to the

southwest. Of the 7 contacts between the Steep Rock Group and the Tonalite

Complex, 3 expose the unconformity (The Headland, S. Roberts Pit, Trueman

Point), and 4 are faulted.

At the Headland poorly sorted metaconglomerate with angular clasts of

quartz, tonalite and fine grained mafic material (dykes and remnant

xenoliths) overlies mafic tonalite, with no evidence for a fault or an

intrusive contact.

At the S. Roberts Pit, poorly sorted metasandstones dip steeply to the west

overlying pale greenish white weathered mafic tonalite. The metasandstones

pass upwards within 20 cm to massive dark grey carbonate.

At Trueman Point, in an exposure similar to the S. Roberts Pit, coarse

angular metasandstone overlies tonalite. However, the contact here is more

diffuse with the top metre of the tonalite breaking down to form a regolith

of angular quartz grains (l-4mm) in a sericite matrix. This matrix is

similar to the overlying metasandstone.

These three outcrops demonstrate unequivocally that the Steep Rock Group was

laid down unconformably on the underlying Tonalite Complex, which is circa 3

Ga old (Davis, pers. comm.).

Contact at Trueman Point

-Poorly sorted

•metasandstone

fl—l-Quartz

Sericite matrix

Unconformity

Weathered Tonalite
THE STRATIGRAPHY OF THE STEEP ROCK GROUP 133

Wilks, M.E. and Nisbet, E.G.

Overlying the Basal Conglomerate (0-150m) is the Carbonate Member (0-500m)

throughout which stromatolites extensively occur.

The carbonate is a laminated dark bluish-grey massive rock with major zones

of breccia developed close to fault zones and dykes, which are thought to be

feeders for the overlying volcanics.

From a study of 11 good stromatolitic outcrops a crude stratigraphy of the

carbonate can be set up. Small scale stromatolites occur throughout the

unit, but are best developed near the base. Here simple Stratifera-like

stratiform structures having flat to undulatory laminae develop into

pseudocolumnar laterally-linked structures. These Irregularia-like

structures pass upwards into hemispherical laterally linked stromatolites.

Laminae are wavy .5-3.5mm, and the structures are 5-15 cm high and in basal

diameter. In places branching walled and unwalled columnar forms occur,

with height up to 20 cm.

In the upper part of the Carbonate giant domal stromatolites occur. These

range from domed structures typically about 3m in diameter to tabular bodies

up to 5m or more long and .75m in stratigraphic height.

Near the top of the unit, small mamillose stromatolites form an egg box

fabric with diameters up to 4cm and heights of 1.5cm. Overlying the

Carbonate Member is the Ore Zone which Jolliffe1 divided into a lower Mn

Paint Rock and an upper Goethite Member.

The Mn Paint Rock (3%-18% Mn) is an earthy material with a rude varicoloured

banding, made up of lumps of goethite, hematite, quartz and chert in a

groundmass of the same minerals with calcite, kaolinite, pyrolusite and

gibbsite. The contact with the underlying carbonate is extremely irregular,

with pinnacles of carbonate protruding into the Paint Rock. This contact

has been interpreted as an ancient karst surface by Jolliffel.

The Mn Paint passes sharply upwards into the Goethite Member (Mn < .3%)

which is predominantly brecciated lump ore of goethite (67%) and haematite

(21%) with quartz and kaolinite.

Within the Ore Zone thin layers of Buckshot Ore occur. These layers

comprise haematitic pisolites and fragments of haematite set in a lighter

aluminous matrix of kaolinite and gibbsite. This material resembles a

ferruginous bauxite in both outward appearance and chemical and mineral

content.

Overlying the Ore Zone is the Ashrock. The name refers to an ultramafic

pyroclastic rock (22% MgO) which makes up to 90% of the unit. Interbedded

within this are thin komatiitic basalt (15% MgO) spinifex-textured lava

flows.
THE STRATIGRAPHY OF THE STEEP ROCK GROUP

134

Wilks, M.E. and Nisbet, E.G.

Within the Goethite Member and Ashrock, pyrite lenses occur. These form

discontinuous elongate bodies of massive pyrite closely associated with

cherty and carbonaceous beds.

The Ashrock delineates the upper member of the Steep Rock Group.

References:

1 Jolliffe, A.W. (1955) in Econ. Geol., vol.50, p. 373-398.
AUTHOR INDEX

 135

Abbott 0. 1

Allen A. R. 64

Attoh K. 4,120

Ayres L. D. 126

Balakrishnan S. 52

Barton, J. M., Jr. 23,129

Bauer R. L. 7,62

Begin N. J. 118

Bhaskar Rao Y. J. 9

Butler J. C. 11

Byerly G. R. 14,72

Card K. D. 17

Chutter G. M. 64

Corfu F. 20

Davis 0. W. 20

De Beer J. H. 23,129

De Wit M. J. 25,27,30

Orury M. J. 33

Elthon D. - 35

Embleton B. J. J. 100

Farquhar R. M. 114

Francis D. 108

Fripp R. E. P. 39

Gibb R. A. 123

Gibbs A. K. 40,85

Glikson A. Y. 44

Grieve R. A. F. 123

Groenewald P. B. 64

Gupta V. K. 123

Hanson G. N. 52

Harper G. D. 47

Hart Rodger 27

Hart Roger 27

Hausel W. D. 50

Helmstaedt H. 55

Hoffman P. F. 118

Hoffman S. 1

Hubert C. 59

Hudleston P. J. 7,62

Hunter 0. R. 64

Hynes A. 108

Krogh T. E. 20

Krogstad E. J. 52

Kusky T. M. 69

LaBerge G. L. 102

Liu M. 108

Losier L. 123

Lowe D. R. 14,72

Lucas S. B. 118

Ludden J. N. 59

Maczuga D. 94
136

Martyn J. E. 75

McGill G. E. 78,106

Morgan P. 81

Morrison D. A. 94

Naqvi S. M. 83

Nisbst E. G. 132

Olszewski W. J., Jr. 85

Padgham W. A. 55

Paris I. 89

Percival J. A. 91

Phinney W. C. 94

Poulsen K. H. 97

Preston V. A. 64

Rajamani V. . 52

Rivard B. 108

Roerlng C. 129

Schmidt P. W. 100

Schultz-Ela D. 62

Schulz K. J. 102,120

Schwerdtner W. M. 105

Scott D. J. 118

Shrady C. H. 78,106

Skulski T. 108

Slejgh D. W. W. 64

Smit C. A. 111,129

Smith P. E. 114

Smith R. G. 64

Southwick D. L. 7,62

Srinivasan R. 117

Stamatelopoulou-Seymour K. 108

St-Onge M. R. 118

Stettler E. H. 23,129

Sullivan R. W. 91

Sylvester P. J. 120

Tatsumoto M. 114

Thomas M. D. 123

Thurston P. C. 123,126

van Reenen D. D. 23,129

van Schalkwyk J. C. 129

Vearncombe J. R. 23,111

Versfeld J. A. 64

Wilks M. E. 132

Wilson A. H. 30,64

Wirth K. R. 86

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