3313:::__:_::_:___;3;:73222.2: mm H

 

This is to certify that the

thesis entitled

GEOCHEMISTRY OF THE HEMLOCK METABASALT
AND KIERNAN SILLS, IRON COUNTY, MICHIGAN

presented by

Thomas P. Fox

has been accepted towards fulfillment
of the requirements for

Ma ste rs degree in Geo logy

 

 

/_ ,
UL. I [LL/III ITLK xf/

Major professor

 

 

Date January 7L1983

 

 

———f'

0-7639 MS U is an Affirmative Action/Equal Opportunity Institution
’W .._-____._- A...» Mfi’}
, - ~ . *3? V
I LE .3ng a is t
I - iii? 4} -
, Micki T3; is case
L University I

 

-v'

 

 

MSU

LIBRARIES
“

 

 

RETURNING MATERIALS:
Place in book drop to
remove this checkout from
your record. FINES will
be charged if book is
returned after the date
stamped below.

 

 

 

 

 

GEOCHEMISTRY OF THE HEMLOCK METABASALT
AND KIERNAN SILLS, IRON COUNTY, MICHIGAN

By

Thomas P. Fox

A THESIS

Submitted to
Michigan State University
in partial fulfillment of the requirements
for the degree of

MASTER OF SCIENCE

Department of Geology

1983

ABSTRACT

GEOCHEMISTRY OF THE HEMLOCK METABASALT
AND KIERNAN SILLS, IRON COUNTY, MICHIGAN

By

Thomas P. Fox

The Middle Precambrian Penokean Orogeny has been attributed to plate
margin processes, to plate rifting, and to formation of an intracratonic basin. In
the Kiernan Sills area, no siallc basement has been identified and good exposures
of Baraga Group mafic volcanics and intrusions are found. The Baraga Group is a
thick sequence of pillow basalt, turbidites, cherts, and pyroclastics structurally
similar to the upper portion of modern oceanic crust. A geochemical study may
help reveal the tectonic origin of this assemblage.

This investigation confirms that the Hemlock metabasalt and west Kiernan
sill are tholeiitic and probably comagmatic. Tentative modelling indicates
fractionation of theoretical parent materials can account for the variations in
basalt chemistry. Tectonic discrimination diagrams give ambiguous results but
generally indicate an extensional environment existed. No conclusive evidence

has been found for an oceanic origin of the mafic rocks.

ACKNOWLEDGEMENTS

John Wilband, my research advisor, has earned my gratitude and deserves
much credit for the completion of this research. His guidance and computer
wizardry have been invaluable.

Thanks also go to Dave Larue, who provided an outline of the problem, the
samples, and funding from N.S.F. grant EAR-81-085114. He also took the time to
review much of my initial work, in essence an unofficial committee member.

I also thank Maria for withstanding the ordeal of being a graduate student's

wife with love and a smile.

ii

TABLE OF CONTENTS

LIST OF FIGURES O O O O O O O O O O O O O O O O O O O O O
LI ST OF TABLES O O O O O O O O O O O O O O O O O O O O O O
INTRODUCTION O O O O O O O O O O O O O O O O O O O O O O

RegionalOverview. . . . . . . . . . . . . . .......

PREVIOUS WORK O O O O O O O O O O O O O O O O O O O O O O

Description of the Hemlock Formation . . . . . . . . . .
Description of the Kiernan Sills . . . ..... . . . . .

ANALYTICALRESULTS...................

Major Elements . .

TraceElements...... ..... .
Metabasalts O O O O O O O O O O O O O O O O O O O
Kiernan Sill O O O O O O O O O O O O O O O O O O O

DISCUSSION O O O O O O O O O O O O O O O O O O O O O O O O
Element MObnity O O O O O O O O O O O O O O O O O O O
Some Comparisons to Continental and Oceanic Environments
REE Fractionation During Igneous Processes . . . . . . . .
TestsofTectonicSetting . . . . . . . . . . . . . . . .
Nature of Parental Magma and Variation of Evolved Rocks .

CONCLUSIONS O O O O O O O O O O O O O O O O O O O O O O O

APPENDIX AO O O O O O O O O O O O O O O O O O O O O O O O

APPENDIX B O O O O O O O O O O O O O O O O O O O O O O O O

APPENDIX CO O O O O O O O O O O O O O O O O O O O O O O O O O

APPENDIX DO O O O O O O O O O O O O O O O O O O O O O O O
REFERENCES O O O O O O O O O O O O O O O O O O O O O O O

iii

iv

vi

Figure 1.
Figure 2.
Figure 3.
Figure #a.

Figure #b.
Figure 5a.

Figure 5b.
Figure 6a.
Figure 6b.
Figure 6c.
Figure 6d.
Figure 7.

Figure 8.

Figure 9a.
Figure 9b.
Figure 10.

Figure ll.
Figure 12.
Figure 13.

LIST OF FIGURES

Regional setting ..................

TheKiernanSills...............

Metamorphic zones within northern Michigan .....

Normative nepheline-olivine-quartz diagram
dividing alkaline from subalkaline rocks . . . . .

Al 03 vs. normative plagioclase plot dividing
cagc-alkaline from tholeiitic rocks. . . . . . . .

AFM diagram for this investigation and sill
samples from Bayley (1959a) . . . . . . . . . .

AF M diagram of metabasalts from Cudzilo (1978) .

REE patterns of the Hemlock metabasalts . . . .

REE patterns of the Hemlock granophyres .......

REE patterns of the Hemlock metagabbros . . . .
REE patterns of the Hemlock ultramafics . . . .

Average REE pattern for each suite illustrating
fractionation trend . . . . . . . . . . . . . .

Range of metabasalt-metagabbro overlap . . . .
Ti-Zf diagram 0 o o o o o o o o o o o o o o o
Ti-Zr diagram from Cudzilo (1978) . . . . . . .

TiOz-KZO-P205 diagram dividing oceanic

fromnon-oceamcrocks . . . . . . . . . . . . . .

F‘MgO‘A1203 diagram 0 o o o o o o o o o o o o
Ti‘v diagram 0 o o o o o o o o o o o o o o o o

Hf/Thvs.Ta/Thdiagram. . . . . . . . . . . .

iv

17

18

19
20
2‘!
24
25
25

34
34
40
41

43
44
as
47

Figure 1‘}.
Figure 15.
Figure 16.

FigUre 17.
Figure 18.

LIST OF FIGURES (continued)

Hf/B‘Ta‘Th diagram 0 o o o o o o o o o o o o o

Mg-Value VS. Th plOt. o o o o o o o o o o o o 0

Normalized Ce/Yb vs. Ce ppm modelling
igneoussystems................

Ce/Yb vs. Ce ppm modelling mineral fractionation

Olivine-clinopyroxene-silica diagram

48
52

55
57
59

Table 1.
Table 2.
Table 3.
Table 4.

Table 5.

LIST OF TABLES

General geology of the Kiernan Sills area. . . . .
\

AveragedREEratios . . . . . . . . . . . . .

Primitive equivalents of selected parental magmas

Coefficients of linear regression to generate
parental basaltSO O O O O O O O O O O O O O O O

Partitioning coefficients used in Figure 17 . . . .

vi

10
22
51

5‘}
58

INTRODUCTION

Much of the metamorphism and structural deformation found in the
Precambrian rocks of northern Michigan and northeastern Wisconsin is related to
the Penokean Orogeny (1.85-I.9 Ga; Van Schmus, 1974). The orogeny has been
attributed to intracratonic activity (Morey and Sims, 1976; Sims et al., 1980),
and to plate margin tectonism (Van Schmus, 1976; Cambray, 1978; Larue, 1981).
During an early phase of the orogeny, volcanism and related intrusions occurred
during deposition of the Baraga Group. The Kiernan Sills in Iron County,
Michigan (Figure 1), formed during this event. Although the sills are within a
zone of low-grade metamorphism, they represent the best exposures of mafic
intrusive igneous rocks related to the Penokean event. As such, trace and major
element data from the sills and the Baraga Group volcanics may provide an
insight into the past local tectonic framework and petrogenesis so as to more

clearly define the models proposed for the causes of the Penokean event.

Regional Overview

In northern Michigan, the Marquette Range Supergroup (Cannon and Gair,
1970; Cannon, 1973) consists of early Proterozoic metasediments and
metavolcanics. The sequence is greater than five miles thick and was deposited
unconformably upon two Archean basement types. The northern basement
complex is composed of granites and greenstones; the southern is gneiss (Morey
and Sims, 1976; Sims et al., 1980).

The Marquette Range Supergroup (MRS) consists of three transgressive

sequences: The Chocolay Group, the Menominee Group, and the Baraga Group

AGE in up moEmH E9; boa—woe: 9:38 3533— ._ 233m

a 9.0.:
.0 92¢

.ookzom_

9 11111111 L

2.92 3.533....

B”.

2.2.: 5:253...

D

2.2: 323.582;

20:. M3323

m

 

(youngest). The stratigraphic position of the Paint River Group is equivocal
(James et al., 1968; Dutton, 1971; Cambray, 1978) and it is not considered here.

The Chocolay Group is a shallow epicontinental sea or miogeoclinal
sequence consisting of basal conglomerate, quartzite, and algal dolomite
deposited uncomformably on Archean basement. These strata are thicker in the
south-southeast of northern Michigan and were apparently derived from uplifted
Archean basement to the north (Van Schmus, 1976). They were mildly disturbed
and partly eroded prior to deposition of the Menominee Group.

The Menominee Group lies unconformably upon the Chocolay Group. It is
initially a shallow water sequence of quartzite but is overlain by turbidites and
iron formation. This transition is interpreted as the beginning of the Penokean
Orogeny by Cannon (1973). As a result of this activity, Menominee Group rocks
in Michigan were uplifted and partly eroded prior to Baraga Group deposition.

Rapid subsidence of the area resulted in a deep-water eugeoclinal
environment. The Baraga Group is therefore represented by turbidites and
volcanics, mostly pillow basalts. The distribution of the volcanics (the Hemlock
Formation and the Badwater Formation) is limited to the west of the Mitchigan
River trough (Cannon and Klasner, 1976). Throughout this area they were
unconformably deposited on Archean crystalline basement or on the Chocolay
Group and were warped by gneiss domes. Radioactive dating of the Hemlock
Formation yields an age of 1.9 billion years (Banks and Van Schmus, 1972). The
composition of these volcanics changes from felsic in the northeast to
intermediate and mafic in the south and west (Cannon and Klasner, 1976). The
best outcrops are located on the southwest edge of the Amasa dome. Tholeiitic
magmatism was very active here where a pair of one-mile thick intrusives (the

Kiernan Sills, Figure 2) and a later mafic plutonic stock (the Peavy Pond

 

 

R.32W. R.3|W.
mc
bh m9
. T
. . fl 2
“‘9 hi“ N.

 

 

zuaa

 

 

 

 

 

 

 

 

 

 

 

Figure 2. The Kiernan Sills (generalized from Dutton and Linebaugh, I967).
mc - Margeson Creek Gneiss, cr - Randville Dolomite, bh - Hemlock Formation,
ba - Amasa Formation, bm - Michigamme Slate, bb - Badwater Greenstone,
pr - Paint River Group, mg - metagabbro. A - discordant contact, B - largest
fragment of Hemlock Formation within the sill. Dots are sample locations.

complex) were emplaced into a miles-thick succession of Hemlock pillow lava.
The west Kiernan sill shows gravity stratification of minerals (cumulate rocks).

Following deposition of the Baraga Group, Penokean orogenic activity
increased in Minnesota and Michigan, being most intense in the south and
southeast, and finally peaking about 1.85-1.9 billion years ago (Van Schmus,
1974). The deformation appears not to have involved much compression of the
Archean basement, which was largely deformed by block faulting creating
structural troughs geographically restricted to the southern part of the Lake
Superior region (Cannon, 1973; Van Schmus, 1976; Baragar and Scoates, 1981).
Somewhat anomalous fold trends, such as the Republic trough, were formed as a
result of the stress distribution across troughs with variable initial orientation.
The troughs were centers of high strain as they were compressed, whereas the
intervening platform sediments exhibit less strain (Cambray, 1978).

Metamorphism is typically a low pressure, low- to high-temperature
variety. Mapping of metamorphic isograds has revealed the presence of annular
nodes, which show chlorite zones at the periphery of each node to zones as high
as sillimanite at the center (Figure 3). Metamorphic grade, like the intensity of
deformation, increases to the south.

In northern Wisconsin, an igneous complex consisting of dioritic to granitic
rocks intrudes a complex of essentially contemporaneous volcanic rocks
(Van Schmus, 1975). Van Schmus (1976) suggested these rocks are related to the
Penokean Orogeny, and may represent the basement of a magmatic arc. The arc
is separated at least in places from the MRS by a major east-west trending fault
system (Dutton and Linebaugh, 1967). The close association of differentiated
mafic intrusives within a thick sequence of pillow lavas, chert, iron formation,
and turbidites, in conjunction with the fact that no Archean sialic basement has

ever been identified between the Amasa dome and the arc terrane, has led some

4.9mq
:7. 6 85mm see 8583 5365 $53: 553 8:8 25.8.2»: .n 23mm

on... 2:35...»

.93. 2:235
02... .0500
329 2:20

82.. 2:23

 

to suggest this area as a possible site of preserved Proterozoic oceanic crust.
The sequence closely corresponds to the upper portion of an ophiolite (Menzies
and Moores, 1976).

The causes of the tectonic activity of the Penokean Orogeny have been
variously interpreted from sedimentological, structural, and geophysical work.
The traditional tenet of vertical remobilization of intracratonic basement has
been steadily advanced by Sims (1976) and Morey (I978). The foundering of
continental crust to form an intracratonic basin above and along the unstable
Archean boundary previously mentioned is attributed to differential thermal
movement of the crust by crustal, and perhaps also mantle, processes (Sims et
al., 1980). Plate margin processes are considered to be nonfunctioning or
unimportant at this time (Sims, 1976). Extensional tectonism did not separate
the crust sufficiently to develop oceanic crust within the basin (Sims et al.,
1980).

Following a plate margin tectonic model, Van Schmus (1976) suggested the
southern (present) limit of Archean rock represents the edge of an Andean-type
continental margin, bordered by the arc complex in Wisconsin. The deformed
MRS sediments north of the plutonic complex represent a deformed back-arc or
foreland basin sequence with subduction occurring to the north beneath the basin
and arc. Upon discovery of Archean basement in Wisconsin, he (1977) later was
compelled to abandon the model.

Cambray (1978) proposes the coincidence of the arc and deformed
sediments is the result of a Proterozoic continental collision. In his model the
Penokean Orogeny was caused by the collision of a northern continent bearing
the MRS on a passive margin miogeoclinal platform with a southern continent
composed of Archean basement intruded by the calc-alkaline magmatic arc.

This collision was the culmination of an intraplate rifting event (analagous to the

modern Red Sea) represented by the MRS and associated volcanics; the entire
sequence of events is the Penokean Orogeny.

One key to the problem of which model should be favored is whether
oceanic crust was produced between the Proterozoic plates during the rifting
event and has been subsequently preserved near the plate suture; or whether the
rifting represents intracratonic crustal foundering of blocks composed of a
voluminous pile of sediments and volcanics above the sialic basement, not
progressing to the point of oceanic crust formation within the basin. As
mentioned previously, the only area where sialic basement is not exposed (or is
missing "?") is between the Amasa dome and the arc terrane. Surficial mapping
and geophysical investigations have not provided a clear picture because of a
lack of exposure in the former case, and the latter do not clearly define the
existence of a sialic basement. Therefore, a comprehensive geochemical study
has been undertaken with D. Larue in the southern Lake Superior region
(N.S.F. grant EAR-81-O85111); this investigation, which focuses on the Hemlock
Formation and intrusive Kiernan Sills, represents part of this ongoing project.

The geochemistry of the rocks may provide information as to their tectonic
affinities to modern environments and thereby provide supportive evidence for
either the plate rifting model or intracratonic basin model as a cause of the

Penokean Orogeny.

PREVIOUS WORK

Description of the Hemlock Formation

The Hemlock Formation is composed mainly of altered basalt flows with
intercalated pyroclastics and sediments deposited disconformably upon the
Randville Formation (Table 1). The Hemlock Formation reaches a maximum
thickness of 25,000 to 30,000 feet near Amasa, but thins regionally to the north,
south, and east (Bayley, 1959b; Wier, 1967). Where it is thinnest, basalt flows
dominate the lithology. Pyroclastics become more common in a westward
direction (upper portion). Gair and Wier (1956) and Wier (1967) both consider the
westward thickening and concomitant increase in pyroclastics to indicate a
greater proximity to centers of eruption, in agreement with the regional
interpretation of Cannon- and Klasner (1975, 1976). The thickness and increase in
sediment content upward is also comparable to the upper portion of an ophiolite
sequence (Menzies and Moores, 1976).

The metabasalt flows are localized, range from 10 to 50 feet thick, and
commonly have vesicles less than one-half inch across or have ellipsoidal tops.
Pillow sizes vary from several inches to several feet, but are usually one to two
feet in diameter (Wier, 1967). Top directions in the vicinity of the Kiernan Sills
are consistently to the southwest.

Bayley (1959b) and Cudzilo (1978) have determined the metabasalts to be
tholeiitic. Major secondary minerals present are hornblende (from pyroxene),
Na-plagioclase (from Ca-plagioclase), chlorite, clinozoisite, and epidote. The
original basalt appears to have had a fine- to medium-grained subdiabasic
texture. In view of the texture and probable metamorphic origin of the

Na-plagioclase, Bayley concludes that the basalts are not spilitic.

10

 

 

30:0 x080 cowowamfi

:mtnEmuoam .8304

 

 

aancoEoocD
035200 macro—Ba

@320 $3850

 

 

>3E3~=oucD

0:5ch cutvoou
c2583"— x0250:
coflmEcom ammE<

azficoHcouca

 

 

mum—m oEEmHfixE
0:39.080 35323

9.8333
@380 owns—mm owcmm
3.33—6.32

 

A5538: 8535
99.0 avid Emma—

 

 

Afimtouc: 538$
Amam 5:..on— umoB cam “35 oEnmwfloE

onEoU vcom broom
zqucoucouca

 

53$qu 2622

 

 

.35 8330

323 539.0030!

earn—p.895 coda:

 

 

 

 

xxficficooca
oxEo—oa can ocean—6:3 :mtnEmU coda: 53.568
ant—53:00::

332.3 BUM—U ocouoafloi .CmEonao

 

 

432 :30 Ba 858 52 .zmamnocs us. 8:5

Eat oo—EEouv moan mam 553v. 05 no 322% 3.050 ._ 03m...

11

No sediments or pyroclastics have been investigated in this research,
however some background on them is useful.

The pyroclastics are typically tuff or breccia. Variations in thickness have
been attributed to post-deposition localization in sedimentary basins. Wier
(1967) has noted graded beds of pyroclastics becoming younger to the southwest,
in agreement with the pillow basalt tops.

Other lithologies in the Hemlock Formation include dark slate and minor

iron formation.

Description of the Kiernan Sills

 

The two Kiernan Sills were first mapped and described, at least in part, and
named by Gair and Wier (1956). Outcrops are irregular and rugged, often
surrounded by marshy lowland. Excellent general and geological information is
available in Gair and Wier (1956), Bayley (1959a, 1959b), and Wier (1967); a brief
summary is presented here.

The sills generally follow the southeast to south trend of the intruded
Hemlock Formation, composed principally of ellipsoidal metabasalt (Figure 2)-
The west sill is the larger and most studied of the two. It is approximately
12 miles long, dips 70°SW to 900W concordant with the Hemlock Formation (Gair
and Wier, 1967), and has a mapped width of 2000 to 6800 feet, averaging 5000
feet (Bayley, 1959a, 1959b). Differentiation by mineral fractionation and gravity
sorting during crystallization (Gair and Wier, 1956; Bayley, 1959a) has resulted in
five distinct zones (Bayley, 1959b): 1) a basal ultramafic zone of metaperidotite,
2) a zone of normal or ophitic metagabbro, 3) a local iron-rich transition zone,
(5) local pockets of granophyre, and 5) a local zone of metadiabase (chill).

The basal ultramafic zone is found intermittently along the eastern portion
of the sill, varying from 600 to 1200 feet thick. It consists of serpentinized

metaperidotite and picrite. Hand samples characteristically show dull green

12

altered phenocrysts of olivine or pyroxene up to one-half inch in diameter in a
darker, greenish-gray fine-grained matrix. Layering less than one inch is often
clearly visible. Other minerals include tremolite, talc, magnetite, carbonate,
chlorite, and actinolite, all probably secondary (Bayley, 1959a; Wier, 1967).

The metagabbroic zone composes the bulk of the sill. Bayley (1959a)
reports that a gradation down to the ultramafic zone and up to the transition
zone can be shown in the southern portion of the sill, best exemplified by
increasing amount and crystal size of original plagioclase and decreasing mafic
minerals up the sill. He also reports rhythmic layering up to six inches is visible,
but only layering less than one-half inch was located during this investigation.

Minerals present (primarily secondary) are albite, oligoclase, actinolite,
tremolite, chlorite, serpentine, epidote, clinozoisite, calcite, stilpnomelane,
ilmenite, sphene, magnetite, quartz, and K-feldspar. Remnant calcite
plagioclase and pyroxene are locally found in the lower metagabbro, but they are
usually replaced by Na-plagioclase, amphibole, and chlorite. Relict diabasic,
gabbroic, and ophitic textures are common (Wier, 1967).

The transition zone is located locally between the metagabbro and the
granophyre, only where the sill is concordant with the Hemlock Formation. The
zone is noted as being gradational and about 200 feet thick in the southern part
of the sill, but Wier (1967) found the zone to be narrow and poorly defined in the
north.

In hand specimen transition rocks are dark brown, fine grained, and
weather to a coarse rubble.

Transition rocks are discriminated from metagabbro by having abundant
ferric stilpnomelane, moderately abundant apatite, more ilmenite and magnetite,

darker amphibole, moderate amounts of quartz, and tourmaline. Often the

13

amphiboles have been replaced by stilpnomelane, chlorite, and epidote (Bayley,
1959a%

The granophyric zone is located near the top of the sill but only in local,
discontinuous pockets. It is hard to differentiate the granophyre in hand
specimen; this investigation has revealed that several outcrops previously
mapped as granophyre are in fact more gabbroic.

The zone is easily characterized in thin section by an abundance of
Na-plagioclase, K-feldspar (often twinned), free quartz, and micrographic
intergrowth of quartz and feldspar. Bayley (1959b) noted the close resemblance
of the Kiernan granophyre with that of other differentiated sills.

Secondary and accessory minerals include chlorite, sericite, biotite,
epidote, carbonate, apatite, sphene, magnetite, and zircon. Feldspars are usually
cloudy and calcite (ferruginous) is pervasive (Bayley, 1959a).

The metadiabasic (chilled) zone is found locally within 100 to l400 feet of
the Hemlock Formation roof. It is not clearly seen in the northern part (Wier,
1967). Texturally and mineralogically it is similar to the metagabbro and
metabasalt but it possesses an intermediate grain size. The typical metamorphic
mineral assemblage is albite-actinolite-epidote-chlorite (Bayley, 1959a).

Pillow basalts adjacent to the sill showing tops facing west, bedding in
Hemlock Formation sediments younging west, and mafic minerals concentrated
on the east (bottom) side of the west sill have all led Bayley (19593) to conclude
that the sill differentiated while cooling in a horizontal position and was
subsequently folded to its present position. A 5000 foot discordant contact
between the sill and Hemlock Formation (Figure 2) indicates the intrusion
occurred at shallow depth and domed at least this much overlying rock.

Portions of the Hemlock Formation can be found as xenoliths in the sill.

The largest fragment is located in the southern portion of the sill between the

1‘}

ultramafic and metagabbroic zones (Figure 2); its displacement is about 3000
feet from its original position at the west (top) of the sill (Bayley, 1959a).
Assimilation of Hemlock Formation appears to be minor, except along the roof
contact which is very gradational and homogenized (Wier, 1967).

Both sills are completely within the greenschist facies of regional
metamorphism (Figure 3), as are all of the Hemlock metabasalt samples. The
fact that primary igneous textures prevail in the metagabbro indicates that the
alteration is essentially low-grade hydrothermal. Heat and water were of
sufficient quantity to transform all amenable rocks to the final stable
assemblage isochemically (Gair and Wier, 1956; Bayley, 1959b). The existence of
local areas which still contain original labradorite and pyroxene indicates that
fluids have been unequally distributed during metamorphism, probably due to the
low permeability of the sill and its great thickness (Bayley, 1959a). Effects of
deuteric and regional metamorphism cannot be distinguished, but Bayley
concludes that most of the chemical differences are due to differentiation of the
gabbroic magma and late deuteric alteration. Regional metamorphism was most
effective on the metagabbro and metadiabase, slightly effective on the
ultramafic and transitional zones, and had no effect on the granophyre.

The east sill is only about half the size of its counterpart, with a maximum
thickness of l#200 feet (Bayley, 1959b) and a length of six miles. No
differentiation is evident; the gabbroic body apparently crystallized in place.
Why only one of two essentially parallel sills of comparable thickness and
original composition has differentiated is uncertain. It is possible that the east
"sill" is in fact a dike (Bayley, 1959b); or that the east sill is perhaps younger,
heating the surrounding rock sufficiently to force the later west sill to cool

slowly and differentiate.

15

No samples were collected from the east sill, although this information
could prove useful in determining if the sills are connected at depth as indicated

by Gair and Wier (1956; plate one cross-section).

ANALYTICAL RESULTS

Major Elements

Eight samples of the Hemlock metabasalt and eighteen samples of the west
Kiernan Sill were analyzed for major and selected trace elements. Sample
locations are given in Figure 2 and Appendix D; analyses are given in Appendices
B and C.

On the basis of these data, the sill rocks are divided into four types as
follows: five ultramafics, one cumulate metagabbro, ten metagabbros, and two
granophyres (Table 2). The basalt samples include massive flows, amygdaloidal
tops, and pillows.

Both the basalts and sill rocks have tholeiitic affinities based on Irvine and
Baragar's (1971) classification scheme (Figure ‘1). Although this method was
originally devised for volcanic rocks, its divisions are based on composition and it
can be used for plutonic rocks if discretion is used.

The sill rocks have a wide range of compositions from serpentinized
ultramafic cumulates to silica-rich granophyres which display a distinct
tholeiitic differentiation on an AFM diagram (Figure 5). Bayley (1959b) was the
first to chemically describe the sill and recognize its tholeiitic nature. Four
samples from his investigation are included in the AFM diagram.

The similarity of the Hemlock basalt and Kiernan metagabbro in both
major and trace element chemistry can be interpreted to mean the rocks of each
group are comagmatic as suggested by Cannon (1973). Because of this similarity,
sill and basaltic rocks are plotted on the same variation diagrams.

Assuming height within the sill correlates with increasing differentiation,

chemical variation diagrams do not reveal a clear stratigraphic trend for the sill

l6

17

0298

HETBBHSRLT

0713

GRRNOPHYRE ea
HETRGRBBRU

RZSF

ULTRRHHFIC 70

+>DGIBO

 

alkaline

 

 

Figure la. Normative nepheline-olivine-quartz diagram dividing
alkaline from sub-alkaline rocks. The heavy line is the division
proposed by the authors (Irvine and Baragar, 1971).

18

.23— command can 05::
$.09. 03:29: 50.: 05.9.3030 we??? 03.02me 0335.5: .m> mO~_< 8 2335 A3 oSwE

mwmbuoHocgm w>HP¢zmoz
2: om om om om om 0.. mm em 2 o

P b _
+ E
+++

 

+

I

L

9
II

2:66.: By. I 65 mmu

<1
<52?
5%
<1
\
a
\
311
80376

SI

2.29:0 -23

 

 

 

EZ

 

F

O 0293
90 El nETRBHSflLT

I 0719
0 GRHNOPHYRE
80 th I "h 5 HETRGRBBRO

oeu e ‘ 929’
+ ULTRRHRFIC

70 X A x annex (1959!!)
IA
El
80 O mAA
All" \ Rm
so X / A
/ / Ag
40 O / cote-alkaline \
/ \ A
30
20
10
R to 20 30 4o 50 so ‘70 80 so I“

Figure 5a. AFM diagram for this investigation and sill samples from
Bayley (1959a). Division of calc-alkaline and tholeiitic suites from ,
Irvine and Baragar (1971). F = FeO + 0.8998 FeZOB'

20

  

G HETRBRSHLT

 

Figure 5b. Fourteen Hemlock metabasalts from Cudzilo (1978).

21

samples. Determining stratigraphic position for the Hemlock metabasalts is

difficult due to folding, faulting, and discontinuous flows.

Trace Elements

The following elements were analyzed by induction coupled plasma
emission at Barringer Magenta Company, Ltd., Toronto: Be, Cd, Cu, Ni, Ba, Sr,
Zr, V, and Zn. In addition, Hf, Ta, Th, Cr, Sc, Co, and selected rare earth
elements (REE) La, Ce, Sm, Eu, Tb, Yb, and Lu were analyzed by instrumental
neutron activation analysis (INAA) essentially following the method of Gordon et
al. (1968). These data are presented in Appendix C.

Following standard procedure, the REE concentrations were normalized to
chondrite values (Frey et al., 1968) and then plotted logarithmically against
atomic number.

Averaged normalized selected ratios for all rock types are summarized in
Table 2. Averaged REE patterns for the sill suites and the Hemlock metabasalt
are presented in Figure 7 for comparative purposes and to demonstrate the
marked similarity between the basalt and gabbro.

Metabasalts. The chondrite normalized REE patterns for the metabasalts
(Figure 6a) are essentially similar except for the most enriched sample (D71A)
and the LREE depleted sample (D2913). The general pattern has a normal to
slightly depleted Ce value, a LREE enrichment, and a small Eu depletion.
Sample D71A has a noticeable Eu depletion and the highest REE content and
probably represents the most evolved sample of the volcanic pile. Although its
exact position within the pile is difficult to ascertain, it is believed to be one of
the highest (youngest) lavas. The other anomalous metabasalt, 0293, may

represent the least evolved parent material.

22

Table 2. Averaged REE ratios.

 

 

 

La/Sm La/Yb Eu/Eu*
granophyre 2 . 615 2 . 92 0 . 714
metagabbro 2 . 01 4 . 02 1 . 00
cumulate metagabbro 2.19 3.97 1.87
cumulate ultramafic 2.18 7.86 not calculated

(greater than 1.0)

Hemlock metabasalt 2.11 3. 75 0. 90
(minus D2913 and D71A)

 

 

Eu* is the predicted value on a line between Sm and Tb.

23

Kiernan Sill. Samples ASB and A5C possess distinct granophyric texture,
enriched REE patterns (Figure 6b), and a distinct Eu depletion relative to the
metagabbros.

All of the metagabbros analyzed except A29F are similar in REE chemistry
and group together in Figure 6c.

Sample A29F is a cumulate rock and was collected from the only
recognized layered zone of the metagabbro. In outcrop, the layers are less than
one-half inch and are easily distinct only on weathered surfaces. The REE
pattern for this sample is similar to the other metagabbros except for its lower
REE abundance and positive Eu anomaly which may reflect fractionation of
plagioclase into the layer (e.g., Weill and Drake, 1973).

The ultramafic samples do not have typical LREE depleted trends (Frey et
al., 1971). Instead, the analyzed specimens show little variation within the group
and have a LREE enriched pattern and a slight Ce and Eu enrichment compared

to the other rock types (Figure 6d).

24

 

 

 

 

 

 

 

 

 

.2‘
E 4 i
E N:
a - . . _,,
2 - 4“
G ' ‘_-“‘,_-——— v
z. - v ‘--—'
{’24 - _
Z - -- .J
,_, .
a l
g ‘ u
4 r3 nctmsnu
13
LR CE 5" EU TB YB LU
O
303753535051Eisfiisgsnoanvon72
RTOHIC NUflBER
'2
l
1
1
u. 1
'- 4
x 1
g 1
o
z 4
U
\
‘
U
a
C
.2‘
l
j o ”more“:
4
. LR CE 5" EU TB YB LU
a

 

 

 

«6753336033 3360601117372

62.36.1336.
aronxc nunaea

Figure 6. REE patterns of the a) Hemlock metabasalts, and
b) granophyres.

25

 

 

 

 

 

 

 

 

I A astncaaano
.\\
1‘."
E
3
z I
5
BE.
‘
U
S .
K
C
LR CE SP1 EU TB YB LU
:sa i7 so is so it i: is do a. 65 do is to 71 72
HTOflIC NUHBER
'34
+ utrnnnnrrc
E
E
O
z
5’.
3%.
x
U
C
G
2'
d x.
L? CE SH EU YB LU
.6 l 1 - I
- ' ' ' 52 is do .3 as 57 so s: 70 ‘1 72
SO 57 5' 5' so RTOHIC NUflBER

Figure 6 (continued).

REE patterns of the c) metagabbros

(including A29F), and d) ultramafics.

DISCUSSION

Element Mobilit;

Examinations of Archean tholeiite (Condie et al., 1977), regionally
metamorphosed basalt (Muecke et al., 1979; Wood et al., 1976), basalt subjected
to seawater alteration (Ludden and Thompson, 1978, 1979), and metabasalts of
various tectonic settings (Pearce and Cam, 1973; Pearce and Norry, 1979; Floyd
and Winchester, 1975; Winchester and Floyd, 1977; Morrison, 1978) including
ophiolites (Pearce and Cam, 1971; Ferrara et al., 1976; Beccaluva et al., 1977)
have concluded that the elements Ti, P, Zr, V, Hf, Ta, Th, Cr, Ni, Sc, Nb, Y, plus
the HREE's in basalts are resistant to alteration and metamorphism. Other
elements such as Sr, Ca, Ba, K, and Na can be shown to be mobile even during
low-grade alteration however (Humphris and Thompson, 1978a, 1978b).

It is of great significance in this investigation to discuss the effects of low-
grade greenschist metamorphism on the abundance and relative distribution of
the REE's in basalt. Unfortunately, there have been few systematic studies of
REE mobility during metamorphic processes, and results have not been
conclusive. Experimental work on REE mobility has only recently begun
(e.g., Menzies et al., 1979).

Some research has reported REE mobility during submarine weathering and
spilitization. Although the Hemlock metabasalts are not spilitized (Bayley,
1959b), seawater alteration has surely occurred to some trace elements. Ludden
and Thompson (1978, 1979) report that the LREE become enriched during
seawater alteration but that the HREE show no selective mobilization. Hellman
and Henderson (1977) found REE alteration within a single tholeiitic flow grading

into spilite, but noted that virtually no change in the REE ratios was apparent

26

27

between slightly altered and more extensively altered material. Wood et al.
(1976) reported significant LREE mobility in zeolitized lavas from Iceland, but
high correlations of HREE with Ti, P, Zr, Ta, Hf, Y, and Nb indicated these
elements were relatively unaffected. Both Frey et al. (1971:) and Wood et al.
(1976) mentioned alteration produced an irregular REE pattern and would
therefore be recognized. Hellman et al. (1979) attempted to categorize gross
and selective REE mobility and maintained that a systematic REE pattern and
high correlation of the REE with "immobile" elements did not demonstrate the
immobility of all the elements, but rather the results could be due to coherent
mobility. They noted, however, that wide variations occurred in REE ratios
(e.g., La/Sm) and anomalies. Muecke et al. (1979) disputed Hellman's claim: "It
seems unlikely that element migration during metamorphism would preserve such
coherent behavior of elements with different chemical properties" (p. lW9).

Early work by Frey et al. (1968) and Philpotts et al. (1969) discerned no
difference in REE abundances of fresh and altered submarine basalt and
concluded seawater did not affect the REE. In greenschist facies alteration,
several authors concluded that low-grade metamorphism did not significantly
change the REE profile. Herrman et al. (1974) suggested that minor fluctuations
may exist in spilites but in no case could LREE enriched lavas become LREE
depleted by seawater alteration. Smewing and Potts (1976) tested pillow basalts
from the Troodos Massif and found little change in the REE during
metamorphism up to 200°C. Condie et al. (1977) studied the effects of
alteration on Archean tholeiites and concluded that the LREE can be enriched if
more than 1096 carbonization-chloritization occurred but that up to 6096
epidotization had no effect on the REE. Hellman et a1. (1977) demonstrated that
within a single outcrop of basaltic rock within the prehnite-pumpellyite facies

the REE behaved coherently with little LREE fractionation.

28

Humphris et al. (1978) found little change in the REE in variously altered
portions of individual basaltic flows, and they suggested that rock crystallization
history would preferentially control REE mobility more than subsequent
metamorphism. Sun and Nesbitt (1978) advocated that although the REE
patterns, particulary the LREE and Eu, can be affected by metamorphism, the
constancy of the pattern enabled recognition of the primary magmatic pattern.
Muecke et al. (1979) agreed with this, and added that metavolcanics up to
amphibolite and granulite facies showed no REE mobility and strong correlation
with previously mentioned immobile elements. Experimental work by Menzies et
al. (1979) has found the REE to be immobile up to 350°C even when the basalt is
totally altered to clay.

In summary, it appears that the REE are affected by: 1) temperature -
zeolite vs. greenschist facies (Wood et al., 1976; Herrman et al., 19711; Hellman
et al., 1979), 2) length of time of exposure to seawater (Ludden and Thompson,
1978, 1979; Menzies et al., 1977), and 3) rock crystallization history (Frey et al.,
19715; Humphris et al., 1978). Common mechanisms of REE mobility include
adsorption by clay and devitrification (Menzies et al., 1980). Mobilization of the
REE will alter the chondrite pattern and ruin correlations with elements such as
Ti, Zr, P, etc. Also, REE ratios and anomalies will vary greatly in samples taken
from close proximity. Since these effects are not noticeable in the REE results

of this investigation, the patterns are interpreted to be unaltered.

Some Comparisons to Continental and Oceanic Environments

Modern ophiolites are thought to be formed at an ocean ridge, small
oceanic basin, or young island arc, and then either abducted or welded to a
continent during continental collision and closure of the oceanic zone (Upadhyay
and Neale, 1979; Serri, 1981; Montigny et al., 1973; Pearce and Cam, 1971, 1973;

Kay and Senechal, 1976; Smewing and Potts, 1976; Pallister and Knight, 1981;

29

Beccaluva et al., 1977; Venturelli et al., 1979; Ferrara et al., 1976; Phelps, 1980;
Menzies et al., 1980). These processes have also been used to explain
emplacement of the early Paleozoic Bay of Islands ophiolite (Suen et al., 1979)
and two late Proterozoic ophiolites, Bou Azzer (LeBlanc, 1981), and the Avalon
zone (Strong and Dostal, 1978, 1980; Strong et al., 1978).

Modern ophiolites typically show a structural cross-section as follows
(Menzies and Moores, 1976): 1) a layer of deep sea sediments, turbidites, cherts,
tuffs, and limestones occurring above and within the volcanic pile and decreasing
in abundance with depth, 2) a layer of pillowed and massive volcanic lavas which
increase in metamorphic grade with depth, and have been fed from below by a
complex dike system and multiple magma chambers at various stages of
fractionation, 3) a plutonic complex containing mafic dikes and felsic screens, all
fractionated, and 4) mantle peridotite of various lithologies (harzburgite, dunite,
lherzolite).

The most important REE aspects of these layers include: 1) LREE depleted
tholeiites, severely LREE depleted picrites, and occasional LREE enriched
fractionates, 2) LREE depleted feeders, 3) LREE depleted gabbros with positive
and negative Eu anomalies in early cumulates and late fractionates respectively,
a) LREE depleted, Eu enriched cumulate peridotites, and 5) flat or LREE
depleted refractory harzburgites and lherzolites. These aspects are present in
both modern and ancient ophiolites (Menzies and Moores, 1976; Sun et al., 1979;
Schilling, 1971; Frey et al., 1968).

Structure and geochemistry have been important factors in recognizing
ancient ophiolites. Studies of the structure and trace and rare earth element
chemistry of the Bay of Islands, Bou Azzer, and Avalon ophiolites indicate that
since the late Proterozoic, oceanic crust has been unchanged except for minor

structural differences. These variations generally include a wider range of

30

ultramafic composition and structure, a diversity of apparent total thickness,
and a lack of a sheeted dike complex in ancient oceanic crust (Burke and Dewey,
1972). The latter two characteristics are not unusual even in modern ophiolites
however (Montigny et al., 1973).

The lithology, geologic structure, and thickness of the Baraga Group of the
MRS is, at least in part, similar to modern and late Proterozoic oceanic crust,
particulary layers 1 and 2 as described by Menzies and Moores (1976). Total
thickness exceeds several miles of pillow lava, black shale, turbidites, chert and
iron formation, and pyroclastics. The sediments increase in abundance up the
sequence, and at least one differentiated sill, the west Kiernan Sill, is present
within the pillow lava. There is no obvious increase in metamorphic grade with
depth, but subsequent metamorphism has had ample time and opportunity to
overprint the original pattern. The absence of layers 3 and It is of little concern;
they may not be present in the Kiernan Sills area because they have not been
recognized due to poor exposure and inconclusive geophysical data.

The LREE enrichment of the metabasalts (and metagabbros) is more
characteristic of modern continental plateau tholeiitic rocks (Schilling, 1971;
Leeman, 1977) or the Keweenawan diabases and basalts which were emplaced
during early rifting stages of the Lake Superior area about 1.0 billion years ago
(Wilband and Wasuwanich, 1980). The fact that MORE-like REE characteristics
(Schilling, 1971; Sun et al., 1979) are lacking, even though turbidites and pillow
lavas indicate a marine environment existed, may be a reflection of the
magmatic processes which gave rise to the parental magma composition and
subsequent lavas, or the composition of the source material.

Unfortunately, the REE cannot be used to assign an oceanic or continental
character to the samples anyway. Leeman (1977) suggested that the different

REE patterns for the environments may be due to different source material.

31

Schilling (1971) attributed the patterns to different rates and modes of magma
transport from the mantle, induced by the different thickness and composition of
crustal material in the two regimes. These modern controls cannot be
extrapolated to the early Proterozoic.

In summary, the REE geochemistry and present geologic knowledge of the
Kiernan Sills area is insufficient to determine whether its Proterozoic tectonic

environment was intracontinental or oceanic.

REE Fractionation During Igneous Processes

Primary mineralogical support for a fractional crystallization model of the
sill consists of increasing original plagioclase and decreasing mafic minerals up
the sill sequence, and layering. Crystallization was one of the main controls on
the sill's final chemistry.

The REE patterns resulting from this investigation (Figure 7) compare
favorably with patterns from other layered intrusions, and provide further
geochemical support for a fractional crystallization model for the sill. Because
the REE are incompatible during the solidification process, they tend to be
enriched in residual liquids. Figure 7 reveals a progressive increase in REE
abundance best noted by the HREE. Helmke and Haskin (1973) reported an
increase in REE enrichment in successive basalt flows from a single parent.
Haskin and Haskin (1968) studied the Skaergaard intrusion and observed an
increase in the REE concentration with little change in relative abundances. No
sympathetic variation with any major element was seen. Kuo and Crockett
(1979) studied the Sudbury nickel irruptive. They found the chilled border norites
best represented the initial liquid REE characteristics, and although some mixing
of a siliceous component occurred with the basaltic melt, the REE patterns still
showed clear fractional crystallization enrichment. The Skaergaard, Muskox,

Stillwater, and Bushveld REE patterns are included for comparison and are

32

similar to Sudbury. Potts and Condie (1971) have examined a layered ultramafic
intrusion and have found REE pattern enrichment.

Eu is preferentially incorporated into plagioclase (Philpotts and Schnetzler,
1968; Weill and Drake, 1973; Schilling, 1971), and because plagioclase is typically
one of the first formed minerals in a basaltic melt, an excess of Eu is expected
in cumulates and an increasing lack of Eu in successive fractionates, assuming
environmental effects on Eu such as f02 and pressure are favorable. This excess
and lack of Eu corresponds to positive and negative Eu anomalies (or high and
low Eu/Eu*), respectively.

Figure 7 illustrates that Eu is enriched in the ultramafics and A29F, and is
successively less enriched in fractionate compositions. Examination of thin
sections reveals olivine (now serpentine) was the dominant crystallizing phase,
and some relict pyroxene remains. No relict plagioclase is visible in the sampled
ultramafics but it is reported to be present by Bayley (1959a); A29F,
accumulated above the ultramafic zone, has visible relict plagioclase. It must be
concluded that the ultramafics originally contained plagioclase but it has since
been altered. The fact that an Eu anomaly is still present reinforces arguments
for isochemical alteration by Gair and Wier (1956) and Bayley (1959b).

The ultramafic suite is unusually enriched in the REE and LREE compared
to similar cumulates from fractionated intrusions. Results similar to the sill
metaperidotite (Bayley, 1959b) were obtained by Morris (1977) for Keweenawan
peridotite in northern Michigan. He concluded the peridotite was generated by
localized melting of a volatile-rich mantle below a developing rift. This model
seems inappropriate because it implies the sill was derived from two different
intrusions, a gabbroic and ultramafic phase, when in fact the sill ultramafic is a

cumulate.

33

No satisfactory mechanisms based on REE modelling can be offered to
explain the high REE abundances in the ultramafic layer.

The ultramafics also have a marked excess of Ce, while the other samples
show a slight negative to non-existent anomaly. Negative Ce anomalies are
usually traced to seawater altertion of the sample (Ludden and Thompson, 1979;
Robertson and Fleet, 1976) and the extent of the anomaly has been linked to
manganese nodule formation (Masuda and Nagasawa, 1975). However, no
correlation is found between Ce and Mn, and seawater or any other fluid
alteration is discounted because: 1) the basalts would be expected to show more
Ce depletion than the intrusives and this is not found, 2) the ultramafics have a
positive Ce anomaly while the other samples do not, 3) the REE patterns are not
irregular, 1+) relict plagioclase, pyroxene, and igneous texture has been found
locally in the basal metagabbro (Wier, 1967), and 5) the sill is over one mile thick
and is impermeable (Bayley, 1959a). Certainly if some late hydrothermal fluid
could percolate through the sill and surrounding country rock extensively enough
to alter the Ce contents, the ultramafics would have been affected, relict
minerals and texture would have been destroyed, and the chemistry would have
been disrupted (including Eu). Isochemical metamorphism appears to have
occurred.

The ultramafics must therefore be enriched in Ce due to some primary
magmatic condition such as oxygen fugacity or composition that existed during
and/or affected only this suite.

It may be noted in Figure 7 that the LREE portion of the ultramafic
pattern is higher than A29F. The crossing of patterns is not unusual for
cumulates because they do not directly represent liquid composition (Pallister

and Knight, 1981, p. 2681!; Kay and Senechal, 1976).

34

 

I! ICTmT
0 WHYRE
‘ ”£1900“
‘ 829'

4- MIMIC

 

ROCK/CNONDRITE

 

 

  

4 7
_ Ln CE an EU TB YB LU

‘ufihhiitaéo‘ofiho‘ihéiofin
atonlc «unaen

 

 

 

 

 

 

ROCK/CNONDRITE

LR CE SH EU TB ‘ YB LU

airinfidoo'tc'zirsfisihfifinfiu
HTOHIC RUBBER

 

 

 

Figure 7. Average REE pattern for each sill suite to
illustrate the fractionation trend. Sample A29F is plotted
separately. The metabasalt average (without D29B or D71A)
is similar to the metagabbro suite.

Figure 8. Range of metabasalt-metagabbro overlap (D298,
D7lA, A29F removed)

35

The average metabasalt pattern (excluding D29B and D71A) is included in
Figure 7 for comparison to the metagabbro average. La/Sm, Lale, and Eu/Eu*
are very similar between the groups. Figure 8 compares the range of the two
suites and shows that considerable overlap occurs. The two suites are therefore

considered to be comagmatic since they are so chemically similar.

Tests of Tectonic Setting

The immobility of the aforementioned trace elements (i.e., Ti, Zr, etc.)
makes them useful as geochemical discriminants to determine tectonic
environment. Several diagrammatic methods developed for Mesozoic and
Cenozoic volcanic suites are presented here in an attempt to compare modern
tectonic environments to the Proterozoic Hemlock metabasalts and the Kiernan
Sills. Although only basalts are usually used in these methods, the gabbros have
been included in the data because of their chemical similarity. Sample D31D has
been omitted. Samples D29B, D7lA, and A29F have been given special symbols
for their identification.

Because these diagrams are best used when the sample population is large,
the reader is reminded that interpretation of the diagrams is tentative,
especially when the number of samples must be further reduced to meet
compositional limitations suggested by the various authors. Also, some diagrams
are derived, in large part, empirically without specific reference to the nature of
magmatic processes operating within the various tectonic regimes.

Tectonic discrimination diagrams for modern rocks and environments have
met with some success (Wood et al., 1979; Wood, 1980; Bhat et al., 1981; Pearce
and Cam, 1973; Pearce et al., 1975, 1977; Smith and Smith, 1976; Ferrara et al.,
1976; Noiret et al., 1981; Shervais, 1982). Morrison (1978) questioned the use of
such tectonic diagrams when applications to Skye and Mull metabasalts failed.

In reply, Wood et al. (1979) and Wood (1980) defended their use, stating that

36

Hf-Ta-Th diagrams indicated the Hebridean lavas had been contaminated, and
that when alteration effects were taken into account Morrison's (1978) criticisms
were not justified.

It has been argued that the REE are not useful as tectonic discriminants
(Kay and Senechal, 1976; Kay and Hubbard, 1978; Menzies et al., 1980) because
patterns for various environments, especially oceanic ridge, marginal basin, and
young island arc, are very similar. Patterns of modern continental and oceanic
tholeiites are easily contrasted, but patterns of ancient analogs are unknown
(Schilling, 1971; Leeman, 1977).

The use of modern tectonic discrimination diagrams to diagnose the
tectonic environment of older rocks is controversial. Several authors (Pearce
and Cam, 1973; Pearce et al., 1975, 1977; Wood, 1980) have attempted to use
diagrams to indicate modern tectonic analogs for Archean rocks. Gill (1979,
p. 431) attacked this practice: "...no universally satisfactory analog of the
Archean tholeiite association exists among modern low—K tholeiites of any type.
It is therefore inappropriate to seek information on the tectonic environment of
greenstone belt volcanics by the use of geochemical discriminants developed
from compilations of analyses of Phanerozoic basalts from known tectonic
associations". However, with respect to the Proterozoic, Gibb et al. (1980,
p.152) stated: "Important contributions could be made by comparing the
essential characteristics of Phanerozoic collision orogenies with suspected
Precambrian examples. In our opinion uniformitarian arguments in favor of plate
tectonic processes operating during the Proterozoic appear to outweigh
arguments to the contrary".

Modern plate systems appear to have first developed 2 billion years ago and
have evolved since then (Burke and Dewey, 1972). Hoffman (1973) suggested the

stratigraphic and structural evolution of the Coronation geosyncline (2 Ga) in

37

northern Canada is sufficiently similar to late Phanerozoic geosynclines to
conclude that it is located along a Proterozoic plate boundary. Likewise the
Labrador trough (2 Ga) has recognizable mio- and eugeosynclinal facets thought
to be located on a Precambrian suture line (Windley, 1973). Coincidentally, the
Hemlock metabasalts have been dated at 1.9 billion years (Banks and
Van Schmus, 1972), and were also contained in a basin with recognizable
geosynclinal facies (Van Schmus, 1976) centered over an Archean suture zone
(Morey and Sims, 1976).

Baragar and Scoates (1981) postulated that a belt of mid-Proterozoic
sedimentary and volcanic rocks around the Archean Superior Province could have
originated in a plate tectonic regime. The Lake Superior area is considered to be
the southern (present) portion of this rift system, which "expanded at the
expense of the northern part into an opening ocean and subduction was initiated
at the southern edge" (p. 297). These rifting and subduction phases correspond
very well with tectonic diagram results from this investigation indicating the
Hemlock metabasalt and Kiernan Sill metagabbro have been formed in an
intraplate rift or a back arc basin model. Gibb et al. (1980) also argue for the
existence of plate tectonic processes similar to those operating today to explain
Proterozoic sutures in Canada, in agreement with Baragar and Scoates (1981).
Cannon's (1973) hypothesis of vertical tectonics operating during the initial
phases of the Penokean Orogeny corresponds well to block faulting that is
expected to occur during the initial rifting event of Baragar and Scoates' (1981)
model.

Various authors have also tried to indirectly or directly expand plate
tectonics concepts into the Archean, arguing similarities between Archean and
modern mantle chemistry, structure, and processes. Archean studies by Sun and

Nesbitt (1977) support Ringwood's hypothesis that the source material for the

38

earth is a carbonaceous chondrite-like material, and that the Archean mantle
was chemically heterogenous but undepleted. O'Nions and Pankhurst (1978)
explored the broad similarities in major and trace element abundance between
Archean volcanics and Phanerozoic volcanics and concluded the chemical
variation of the earth's mantle in early Archean times was comparable to that at
the present. Sun et al. (1979) added a note of uniformitarianism by suggesting
mantle depletion is due to a continuing process rather than an episodic event in
the Precambrian.

Hart et al. (1970) compared modern and Archean tholeiites, especially their
trace element contents, and concluded Archean rocks were closely related to
modern low-K tholeiites of island arcs. They therefore proposed that the
Archean tectonic environment was dominated by multiple thin oceanic plates
with arc-trench boundaries and shallow subduction zones. Condie and Baragar
(1974) studied REE patterns of Archean greenstone and noted that replenishment
of source material was required to produce the REE patterns. They suggested
the replenishment mechanism could be subduction, but they cautioned use of
Phanerozoic plate tectonic analogies. Jahn et al. (1974) compared Archean
volcanics from northeast Minnesota to modern equivalents. Trace elements
suggested to them that the rocks had formed in an ancient island arc-subduction
zone, and that the mantle has been a closed, recycling system for the past
3 billion years. Nesbitt and Sun (1976), while acknowledging modern equivalents
to Archean greenstones are lacking, suggested the closest tectonic analogy to
Archean greenstone formation would be development of oceanic crust within a
rifted continental block. Arth and Hanson (1975), Gill and Bridgwater (1976), and
Gill (1979) have attacked the use of geochemical discrimination in assigning a
modern tectonic setting to Archean rocks precisely because petrologic details

between the two sample populations are different, and even differences between

39

the modern types themselves must be understood more fully. Burke and
Dewey (1972) simply do not accept operation of a plate tectonics system before
2.7 billion years.

In summary, it appears to be generally accepted that some form of plate
tectonics analogous to that operating today was in existence in the geographic
position and during the time of production of the Hemlock Formation and the
Kiernan Sills, and perhaps even before. Therefore, in the spirit of Gibb et al.
(1980), the application of modern tectonic discrimination diagrams to these rocks
may have some contribution to understanding the environment of their
formation, and in outlining possible causes for the Penokean Orogeny.

Pearce and Cann (1973) suggest that intraplate, ocean floor, low-K
tholeiite, and calc-alkaline suites can be discriminated by Ti, Zr, Y, Nb, and Sr
distributions. For altered suites they advise a Ti-Zr plot (Figure 9), and a
Ti-Zr-Sr plot for unaltered rocks. They suggest that these diagrams be used for
data which do not plot in the intraplate field from initial Ti-Zr-Y ternary data.
Because Y data are not available, the figures cannot be used to indicate whether
the rocks have intraplate affinity, but show only that the calc-alkaline field is
unoccupied.

Floyd and Winchester (1975) introduced a series of diagrams to
discriminate magma type (alkalic vs. tholeiitic) and tectonic setting.
Continental and oceanic domains of each type were used to determine the
diagrammatic fields. The authors noted that although the magma types are
easily distinguished, the domains overlapped to an extent that no meaningful
separation could be made to distinguish them.

Data from this investigation fall within the tholeiitic fields but cannot be

used for tectonic discrimination because of the inherent overlap of the diagram.

40

 

 

 

 

 

 

 

“J | a fi'
5 I
a .3 s
an 5 ‘
l I t
. I i
O . .
_ 1 i
1 a a I
*4
.1 1 1
O N
C __ a E!
D A
\ O .
Z d
L
Q.
s—v a ‘
l'-- A D
(D 1 3
1
l ’ A a c
v «1
i A o 0sz
El .‘IETRBRSRLT
o. 1 A I 0719
A HETRGRBBRO
A RZSF
o o I I

5 25 so 75 160 155 :so 135 250 2E5 250
ZR ”

Figure 9a.. Ti-Zr diagram. Fields A + B = low-K tholeiite, C + B =
calc-alkaline, D + B = ocean floor (Pearce and Cam, 1973).

41

 

12 I4 16 18 20

TIIPPl‘l/IOOOI
110

 

Ol-

 

I!) "E IflBflSRL I

 

 

 

’ I I I I j I I I
o 25 so 75 160 125 150 175 200 225 250 275 300
ZR

Figure 9b. Ti-Zr d' ram. Data fo .
Cann, 1973). ‘33 r m Cudz1lo (1978). (Pearce and

42

Pearce et al. (1975) proposed that a ternary TiOz-KZO-P205 diagram
(Figure 10) may discriminate between oceanic and non-oceanic primary basalts
with alkali content less than 20% on the AFM diagram. Only D31D exceeds this
limitation.

The elongate trend may be attributable to: 1) fractionation of the samples,
2) metamorphism, which enriches K in oceanic basalts, and/or 3) continental
rifting, which produces rocks with both oceanic and non-oceanic affinity.

The effect of fractionation of the basaltic parent is indicated by D29B and
D7lA, the least and most evolved basalts according to the REE. In this case,
primitive rocks would give the best estimate of original composition. All of
them plot in the oceanic field.

Fractionation in the sill also results in K enrichment. Assuming a sampling
representative of the gabbroic magma has been collected, the sill average plots
near the group of oceanic basalts. This is expected since the suites are
comagmatic.

Metamorphic effects have probably been significant on K content, but the
authors note, "if a metamorphosed basalt falls within the oceanic field..., it is
very likely of oceanic origin" (p. 423).

The TiOZ-KZO-P205 diagram is therefore very important as an indication
that oceanic-type basalts formed in the southern Lake Superior region in the
Proterozoic which supports case three that some form of continental rifting
proceeded to this point, contrary to Sims' (1980) model. A similar conclusion
partially based on this diagram was reached by Bhat et al. (1981) for Panjal trap
rock in India.

A tectonic discrimination diaigram using F-MgO-Alzo3 has been proposed
by Pearce et al. (1977). F is total iron as FeO. Instead of basalt, this diagram

(Figure 11) uses "basaltic andesites" with 51% to 56% SiO2 anhydrous. Only two

43

T102

       

non- oceanic

20

 

o 0298
El HETRBRSRLT
I 0719
10 A nETflGflBBRO
A HZSF
10 20 30 40 50 80 70 80 00
K 2 0 P 2 05

Figure 10. TiOZ-KZO-P205 ternary diagram for oceanic vs. non—oceanic
basaltic rocks (Pearce et al., 1975).

an

 

 

20
c: nernsasntr
m A nernoaeaao
10 20 so 40 so so 70 so so
M00 RL203

Figure 11. F-MgO-A1203 diagram. A = continental basalt field,
B = ocean floor basalt field (Pearce et al., 1977). F = FeO + 0.8998
Fe203.

45

metabasalts met the criterion as did five Hemlock metabasalts reported by
Cudzilo (1978). These two samples plot in the oceanic field of the
TiOZ-KZO-P205 diagram.

Shervais (1982) suggested using the immobile elements Ti and V for
petrogenetic plots of modern and ophiolitic lavas. The results for this
investigation are ambiguous. The scattering of data in Figure 12 prevents
distinguishing whether the MORB, Columbia River flood basalt, or back-arc basin
environment (not shown) is the best analogy. Shervais notes the contrast in flow
morphology and associated sediments between MORB and continental basalts
makes it unlikely these settings will be confused. The results could be used to
rule out the presence of an island arc assemblage. This field plots between a
Ti/V ratio of 50 and 100.

Figure 13 is a plot of Hf/Th vs. Ta/Th used by Noiret et al. (1981) to
successfully interpret the tectonic setting of the Vourinos ophiolite complex in
Greece. The data cluster mainly in the island arc-subduction zone field. This
result is directly opposed to results from the Ti-V diagram. Noiret et al. offer
an alternative to this tectonic assignment: the samples could fit an
intracontinental rift model with a large degree of contamination by continental
crust (p. 383). This alternative is in agreement with results of the
TiOz-KZO-P205 diagram (Figure 11). Direct evidence of contamination such as
resorbed xenoliths or crystals is lacking however, and evaluation of
contamination is difficult.

Wood et al. (1979) introduced and revised (Wood, 1980) a tectonic
discrimination diagram using Hf-Ta-Th (Figure 14). Most samples from this
investigation plot outside the tectonic fields. According to the authors, this

suggests that the sample series has been contaminated since Hf, Ta, and Th are

46

 

 

 

8
u) 1 2' f \ J
1:
L31 A \
/
$1 / so
/ /
8.
a / 4/ I.
C A
Oh
"’ / ‘3 co A
03
s v /
>N I /
c A
g- D f / loo
8. / /A
~ /
SJ 0 0298
- 1:: 11571135151111
I 0719
. A nctnonaano
a 929?
c I I T I

 

 

o i 4 E 51012141518 20
THPPM/IOUO)

Figure 12. T i-V diagram. Solid line - Columbia River basalt field,
dashed line - MORB field (Shervais, 1982).

47

 

 

 

 

 

 

 

on
[a . o 0298
m nernansnLt
I 0719
(0.. A "5190986110
A nzsr
“3.4
v4 Subdoclion
[I] M
O
o co n A N
is l on)
k/B‘
o

 

 

T T F I I F I I F
-o 0.10 0.20 0.30 0.40 0.50 0.60 0.70 0.80 0.90 1.00
TR/TH

Figure 13. Hf/Th vs. Ta/Th diagram (Noiret et al., 1981).

48

 

HF/B
so
so
70
a
El
o 0293
1:! 11579311391:
I 0719
a astaonoono
. 9st X
40 so so 70 so so

TF1

 

Figure 14. Hf/3-Th-Ta diagram. A = N-type MORB, B = E-type MORB
and tholeiitic in-plate basalts and differentiates, C = alkaline in-plate
basalts and differentiates, D = destructive plate margins and
differentiates (Wood, 1980).

49

very sensitive to crustal contamination. This conclusion is in agreement with the
alternative interpretation of Noiret et al. (1981).

In summary, there is no clear definition of tectonic setting for the
Hemlock metabasalts and intruding Kiernan Sill. In comparison to modern suites
of known tectonic environment, the early Proterozoic rocks under investigation
are most similar to those found in: 1) an intracontinental rift environment in
which the magma became contaminated by crustal material, or 2) magma
generation in a back-arc basin. These two environments can be generalized
together to support an extensional tectonic setting in the Kiernan Sills area at
the approximate time of their emplacement. Sims et al. (1980) and Cambray
(1978) have also suggested an extensional regime in the area in the early
Proterozoic.

The tectonic diagrams cannot clearly differentiate whether the samples

are of oceanic or continental affinity.

Nature of Parental Magma and Variation of Evolved Rocks

 

In this study, several igneous processes may have left an imprint on the
elemental abundances of the gabbros and basalts. These are: 1) Partial melting
of a source resulting in generation of primitive magma, 2) evolution of the
primitive magma to a parental composition for the basalt and sill, and
3) evolution of the parent magma to the sample compositions. Once primitive
and parental compositions are established, it may be possible to explain the
chemical variation within the lava pile and establish comagmatism between the
sill and basalt.

Chemical models presented in this section are usually attempted with a
larger sample population; hence, the results must be regarded as tentative.

Primitive magmas are formed by 2% to 30% partial melting of mantle

sources (peridotite) at high pressure. A common indicator of a primitive magma

50

which has not undergone modification en route to the surface is a Mg-value
(100 Mg/Mg + Feb) of 68 to 72 (Frey et al., 1978). By this criterion none of the
samples are primary.

The selection of possible parents is based upon chemistry and texture of
the samples. Compatible trace element (Co, Ni, Cr) abundances are very
sensitive to fractional crystallization and are high in least evolved parent rocks.
The parents will also have low $102, and the highest MgO and Mg-values.
Parental compositions from intrusions are generally believed to be preserved in
chill margins. Based on these criteria, two basalts (D29B and D34G) and one
gabbro near the sill top (ASP) were selected as possible parent material. The
selection of two basalt parents was made because of the distinct differences in
the REE patterns between the two (i.e., sample D29B is LREE depleted).

Estimation of the composition of the "primitive liquids" by back calculation
to an Mg-value of 68 are given in Table 3. The results of these calculations show
that the parental compositions could have been derived by removal of 8% to 15%
olivine (F087) or up to 30% pyroxene (quoEnsoFs 10)’ or a linear combination of
both.

The abundance of a highly incompatible element such as Th, or moderately
incompatible elements such as La or Ce, shows an inverse relationship to
Mg-value, as demonstrated in Figure 15. It is therefore assumed that
compositional variation of the basalts was largely controlled by removal of Mg
probably by fractional crystallization.

The evolution of the presumed parental basalts to later compositions can
be modelled in a closed magmatic system by multiple linear regression whereby
the parent is derived by summing the coefficients in the equation:

Parent = evolved rock + F080 + pyx + plag (An 55) + 11 + mag.

51

Table 3. Primitive equivalents of selected parental magmas
by back calculation for olivine addition.

 

 

 

 

Basalt Gabbro

D29B D34G A5P
SiO2 49.12 49.38 50.19
A1203 12.27 11.88 15.07
Fe203 1.43 l. 55 1.28
FeO 11.17 11.68 9.42
MgO 13.40 13.88 11.39
CaO 9.11 5.71 9.31
Nazo 2.24 3.13 1.69
K20 0.23 0.63 0.63
T102 0.79 1.93 0.79
P205 0.05 0.12 0.05
MnO 0.16 0.13 0.13
Mg-value 68 .13 67 . 93 68 .17
Olivine added 15% 12% 8%

 

 

52

a:

.5. .m> 039-34 .3 0.5!...

 

 

m m w w v m N a o
b 1— . — h _ _ _ 8
0
8
13
17
O
B

1.7
B S
C...
TO
PS
9

$8 I B
38 a 9
Scmcmfiuz 9 BO 10

88 o
E

 

 

99

#011

53

Alteration of the minerals in the samples necessitated using mineral
compositions typical of basaltic rocks (except plagioclase). Results of the
calculation for the basalts are printed in Table 4 in order of decreasing
Mg-value. The bottom row of Table 4 gives the calculated percentages of
minerals required to form D34G from D29B. The negative value for olivine
reflects the fact that 029B has a lower Mg-value than D34G. In other words, it
is difficult to explain the evolution of D34G from D29B by closed system
fractional crystallization because the negative coefficient requires that olivine
be added to the system, whereas in all other systems it must be subtracted
(i.e., evolved rock = parent-OL-PYX-PLAG-IL-MAG).

It is standard procedure to use plots of incompatible element ratios to
evaluate magmatic processes which may have existed within a magma chamber.
For example, LREE/HREE or LREE/intermediate REE ratios are respectively
used as indices of absolute and relative fractionation of the REE. Changes in
these ratios with respect to another incompatible element can be easily modelled
for fractionation (in an open or closed system, or of a specific mineral) or for
partial melting of a source.

The normalized Ce/Yb vs. Ce ppm plot derived by Wood (1978) for
Icelandic basalts is one of many possible standard plots which can be used to
evaluate magmatic processes. It has the advantage of using elements for which
the analytical data are precise and for which the partitioning coefficients for
minerals in basaltic systems are well known. The internal lines in Wood's
diagram are calculated paths of open system fractionation, closed system
fractionation, and partial melting.

The Ce/Yb vs. Ce plot (Figure 16) has a steep-sloped trend which includes
samples D29B, D68A, and D29C. This trend parallels the partial melting or the

open system fractionation trends derived by Wood. A flexure in the sample trend

su

 

 

 

 

 

 

3.8 $o.~ 34 ”N5 26 3.2 no; 3 GE 9% m2
«an 83 as to R4 12 o.o Gas ”.3
as R: E n.~ 12 6.: RA 2ch s8 m3
in 3.3 ”a as «a a; 3: 25 so“
12 1.3.0 5.0 m; 0.8 .2 EB 33 0.8
52 83 m5 ~.~ MEN ms 92 23v 5%
5: £6 ”5 a; k; ca 8.2 85 8.: can
am: .2353 952 a 05a 53 .6 385 8 Evan

N

 

 

8:83 .35ch 320.com 3 3:39. 9.02 3290
van £8058 no commmocwoc have: no 39658 8 #3 23 3565000 .a vim...

55

.852 683 .83
3:032: vocutuoccmw .. n £53.35 amazon 1 a ”no 95.08 3?ch :33 E25530; coco u m 25:93:08“
come? no 389: 53025 n N + _ 6.839? «EEK 95—260.: Ema oU .m> n>\oU 3338.32 .3 v.53"—

21: m5
om S. 8 om 21 am ON 2 o
_

 

 

 

 

 

 

 

_ _ _ _ _ _ 0
Q
~ 1 I 3
I 3..
$2.0 I ut- /
$8 I T 3 IA
Zcmamfiuz a m nd
88 o U ,1
fl‘ I
1.1.1111 w B 8 N
I n
¢ .7

 

56

occurs near sample D34G, where the remaining basalts all plot on a relatively
flat curve which closely parallels Wood's closed system fractionation curve.
Although there are only eight data points, these trends are interpreted to mean
that two processes must be evoked to explain compositional variations in the
lavas.

Partial melting by itself probably cannot account for the compositional
variations because of the stratigraphic placement of the samples. For example,
if the steep-sloped grouping is due to cogenetic derivation by partial melting, the
melt derived by the least amount of melting would be expected to erupt before,
and therefore lie below, lavas which were derived by greater amounts of melting.
This stratigraphic zoning is not apparent in the lava pile. Variation could also be
explained by mixing primitive liquids with evolved magma (open system), but
limited sampling precludes testing this hypothesis.

If one calculates the Ce/Yb vs. Ce points for a system which fractionally
crystallizes olivine, plagioclase, and clinopyroxene (Figure 17) using the Rahleigh
equation and appropriate partitioning coefficients (Table 5), the steep-sloped
sample trend can be attributed to fractionation dominated by clinopyroxene.
Olivine and plagioclase do little to the Ce/Yb ratio during fractionation, thus
their removal leaves the evolved liquids with a shallow to flat slope. These data
are consistent with the major element modelling in which fractionates of D29B
have high pyroxene coefficients (Table 4) and derivatives of D34G have dominant
olivine and plagioclase coefficients.

The mineral fractionation trend can be demonstrated from the olivine-
clinopyroxene-silica-plagioclase tetrahedron, where the mineral percentages are
calculated following the method of Walker et al. (1979), and Stolper (1980). The
internal points are projected onto the olivine-clinopyroxene-silica plane

(Figure 18). The projections of the three phase boundaries at one atmosphere

57

 

 

 

 

 

 

 

 

CPX I 1
LD‘
”3 055
__‘ 8
CD
>- 07“ o 0sz
\ m nemaasau
L” a 0340
Q E“ I 0719
LD-l
m-
- CPX 2
A'PLAG
O 10 20 30 4O 50 60 70 80

CE PPM

Figure 17. Ce/Yb vs. Ce ppm modelling mineral fractionation using data in
Table 5. D298 is modelled as the original composition.

58

Table 5. Partitioning coefficients used in Figure 17.

 

 

 

01.“) cpx1“) cpxzm PLAGQ)
Ce 0.0033 0.166 0.04 0.062
Yb 0.0202 1.01 0.20 0.009

 

 

References: (l) Frey et al. (1978)

(2) Paster et al. (1974)

59

 

 

CPX
O 0298
E] HETRBRSRLT
90 U 0340
I 0719
so 0 GRHNOPHYRE
A HETRGHBBRO
70 ‘ HSP
+- ULTRRHRFIC
00
so
40
A
AA A OI A
30 "0.0:-
20 E139” h m
a E1 “In.
10 A. ”'
4. ‘T mm- (D
#11:. 42;. r. a. .‘a 7". as ‘1'-
OL mm 3102
Figure 18. Olivine-clinopyroxene-silica diagram (after Walker et al.,

1979).

60

and ten kilobars from Stolper (1980) serve as pressure constraints for the system.
The lavas plot with a distinct trend which leads diametrically away from the
olivine corner, then follows a path which subparallels the one atmosphere
boundary toward $102.

The sill rocks have a diverse distribution which reflects the extreme
variation in liquid compositions due to fractional crystallization. As expected,
cumulates are clustered near the olivine corner and granophyres are located near
the SiO2 corner.

Walker et al. (Figure 3, 1979), shows the mixing of primitive liquids with
more evolved magmas causes rocks to plot on the concave side of the olivine-
clinopyroxene boundary. Experimental data for these curves at pressures
between one atmosphere and ten kilobars are lacking, thus this test cannot be

applied to the lavas.

CONCLUSIONS

Geochemical analysis, tectonic comparison using diagrammatic methods
and limited chemical modelling of samples from the Kiernan Sills and
surrounding Hemlock metabasalt allow the following conclusions: 1) despite the
age and metamorphic grade of the area, geochemical analysis, especially for the
immobile elements (e.g., Ti, Zr, REE), has provided useful information and future
chemical studies are encouraged, 2) the gabbroic sill and Hemlock basalt are
tholeiitic and probably comagmatic, in support of earlier studies, 3) several
outcrops of the sill mapped as granophyre are in fact gabbroic; mapping should
be revised, 4) tectonic discrimination diagrams applied to these Proterozoic
igneous rocks yield ambiguous results in comparison to modern tectonic analogs;
liberal interpretations of these diagrams favor an extensional tectonic regime, in
general agreement with present tectonic models (e.g., Sims et al., 1980;
Cambray, 1978) which differ vastly on details of extensional mechanisms and the
extent of crustal thinning, 5) with the exception of one analysis, the tholeiites do
not have primitive "MORB-like" REE patterns and are not alkalic, thus the LREE
enrichment cannot be attributed to small amounts (less than 10%) of source
melting; if melting was in the 20% to 30% range the source must have been
LREE enriched; the patterns could also be due to evolution of the parental
magma, producing LREE enriched samples, 6) limited modelling indicates the
parent magma could have evolved from a primitive melt by removal of 8% to
15% olivine or up to 30% pyroxene, or a linear combination of both; the modelled

parental compositions evolved in pressures less than 10 kilobars.

61

APPENDIX A

Analytical Methods

Instrumental neutron activation analysis (INAA) was used to determine
whole rock compositions of the rare earths La, Ce, Sm, Eu, Tb, Yb, and Lu, and
of Na, Sc, Cr, Co, Hf, Ta, and Th. The equipment used in this analysis includes:

1. A Triga Mark I reactor,

2. A GeLi gamma-ray detector — a lithium-drifted germanium crystal

held at cryogenic temperature by liquid nitrogen,

3. A multichannel analyzer, signal amplification equipment, and other

associated electronics, and

4. The Cyber 750 computer at Michigan State University.

Powdered (1.00000 gram) rock samples, crushed to pass through a 200 mesh
screen, were placed into polyethylene vials and sealed for irradiation. Gloves
were used to prevent contamination and assure true mass was recorded. In
addition the vials were washed thoroughly before irradiation.

The samples were irradiated in two batches. Batch 1 consisted of basalts,
gabbros, and granophyres plus four USGS standards and four evaporated liquid
standards. Batch 2 consisted of ultramafics plus five USGS standards and four
evaporated liquid standards. Each batch remained in the reactor for six hours at

12 neutrons/cmzlsec. Analysis began one week after irradiation

a flux of 10
following the method of Gordon et al. (1968).
Oxides and other trace elements were analyzed by induction coupled

plasma emission at Barringer Magenta Laboratories, Toronto, Canada.

62

APPENDIX B
CIPW OXIDES AND MINERALS

 

 

 

 

 

Metabasalt .

D29B 1329c D29E D3lD D34G
5102 50.44 49.29 51.88 53.09 50.46
141203 14.11 15.09 14.02 15.37 13.31
Fe203* 1.64 1.82 1.95 1.63 1.74
FeO 10.96 12.16 13.02 10.85 11.58
MgO 8.35 7.98 4.83 5.04 9.89
CaO 10.48 8.95 7.72 6.58 6.39
NaZO 2.58 2.80 3.09 4.62 3.50
K20 .27 .37 .77 .86 .70
1102 .91 1.25 2.25 1.62 2.16
9205 .06 .11 .28 .18 .13
MnO .18 .19 .19 .16 .14
Norms
Q 0.00 0.00 2.52 0.00 0.00
c 0.00 0.00 0.00 0.00 0.00
Or 1.60 2.18 4.54 5.07 4.15
Ab 21.87 23.67 26.11 39.10 29.64
An 26.12 27.54 22.12 18.65 18.51
Ne 0.00 0.00 0.00 0.00 0.00
Wo 10.64 6.75 5.98 5.35 5.15
En 14.97 11.42 12.02 7.86 12.40
Fs 12.65 10.99 18.94 10.13 8.32
Mt 2.38 2.64 2.83 2.36 2.52
11 1.74 2.37 4.27 3.08 4.11
Ap .15 .25 .67 .43 .30
01 7.90 12.20 0.00 7.98 14.90
(Di) 20.96 13.37 12.04 10.70 10.08
(Hy) 17.30 15.80 24.90 12.65 15.80
Plag 54.43 53.78 45.87 32.30 38.45
Mg* 59.95 56.31 42.12 47.72 62.64

 

63

64

 

 

 

 

APPENDIX B (continued)
Metabasalt Granophyre
D65A D68A D71A A5B A5C

$102 49 .89 48 .53 49 .93 66 .06 74 . 56
A1203 14.21 16.53 12.90 10.96 12.14
Fe203* 2.01 1.53 2.23 1.45 .58
FeO 13.42 10.17 14.89 9.70 3.88
MgO 5.76 7.95 3.92 1.82 .45
CaO 8.08 12.19 8.24 3.61 2.33
Nazo 1.90 1.77 2.04 3.51 2.25
K20 2.76 .23 2.31 1.77 3.38
T102 1. 60 . 90 3 . 05 . 97 . 37
P205 .19 .07 .32 .06 .01
MnO .17 .14 .17 .10 .04
Norms

Q 0.00 0.00 1.38 23.55 40.11
C 0.00 0.00 0.00 0.00 .57
Or 16.33 1.38 13.63 10.44 19.97
Ab 16.09 14.97 17.27 29.70 19.03
An 22.07 36.47 19.24 8.92 11.51
Ne 0.00 0.00 0.00 0.00 0.00
Wo 7.00 9.81 8.17 3.57 0.00
En 9.74 14.23 9.77 4.53 1.13
Fs 14.02 11.64 20.78 15.19 6.11
Mt 2.92 2.21 3.24 2.11 .84
11 3.04 1.71 5.79 1.84 .70
Ap .45 .18 .75 .15 .03
01 8.35 7.41 0.00 0.00 0.00
(Di) 14.04 19.31 16.60 7.35 0.00
(Hy) 16.73 16.37 22.13 15.94 7.24
Plag 57.83 70.89 52.69 23.10 37.68
Mg* 45.73 60.54 34.10 26.91 18.70

 

 

65

APPENDIX B (continued)

 

 

 

 

Metagabbro
A532 A5P A61 A27F A2A

5102 48.80 50.98 50.32 49.33 46.75
A1203 16.87 16.28 16.00 14.31 12.24
Fe203* 1.51 1.38 1.63 1.92 2.57
FeO 10.10 9.23 10.85 12.81 17.15
Mgo 7.01 8.53 5.59 5.63 5.06
CaO 11.35 10.05 8.95 9.54 8.67
N320 2.02 1.83 3.79 3.03 2.39
K20 .82 .68 .68 .53 1.27
no2 1.25 .85 1.90 2.63 3.45
13205 .13 .05 .14 .12 .23
MnO .14 .14 .16 .15 .22
Norms

Q 0.00 .28 0.00 0.00 0.00
c 0.00 0.00 0.00 0.00 0.00
01' 4.83 4.01 4.01 3.16 7.48
Ab 17.06 15.51 32.09 25.66 20.27
An 34.57 34.18 24.62 23.85 18.90
Ne 0.00 0.00 0.00 0.00 0.00
W0 8.73 6.39 7.89 9.47 9.44
En 11.48 21.24 5.75 10.48 7.73
Fs 10.17 14.65 6.50 13.38 14.76
Mt 2.20 2.01 2.36 2.79 3.73
11 2.38 1.62 3.60 4.99 6.56
Ap .30 .13 .33 .28 .55
01 8.30 0.00 12.87 5.95 10.61
(Di) 17.24 12.51 15.69 18.93 19.12
(Hy) 13.15 29.77 4.44 14.41 12.82
Flag 66.96 68.79 43.41 48.17 48.26
Mg“ 57.69 64.47 50.31 46.30 36.69

 

66

 

 

 

 

APPEme B (continued)
Metagabbro

A4B A4C A5A D29B A29F
5102 47.10 47.06 52.28 52.12 43.21
141203 14.82 14.40 15.70 14.56 19.26
191220311 1.99 1.83 1.24 2.04 1.41
FeO 13.29 12.20 8.25 13.59 9.39
M30 6.86 8.60 7.09 4.38 14.44
C30 9.72 10.20 10.19 7.05 8.84
N320 2.88 2.46 2.82 2.99 1.86
K20 .35 .48 1.13 .45 1.05
1102 2.66 2.43 1.08 2.35 .39
13205 .18 .16 .09 .28 .04
MnO .16 .18 .12 .19 .12
Norms
Q 0.00 0.00 0.00 4.96 0.00
c 0.00 0.00 0.00 0.00 0.00
Cr 2.06 2.86 6.66 2.68 6.18
Ab 24.36 20.82 23.90 25.27 5.91
An 26.48 26.81 26.84 24.99 41.13
Ne 0.00 0.00 0.00 0.00 5.30
Wo 8.58 9.49 9.65 3.39 1.05
En 6.90 8.46 14.74 10.91 .68
F5 7.54 6.81 10.49 19.74 .30
Mt 2.89 2.65 1.80 2.96 2.04
11 5.05 4.61 2.04 4.47 .74
Ap .43 .38 .23 .68 .08
01 15.73 17.12 3.66 0.00 36.60
(Di) 17.05 18.66 18.90 6.85 2.02
(Hy) 5.96 6.09 15.97 27.18 0.00
Flag 52.08 56.29 52.90 49.72 73.61
Mg* 50.32 58.03 62.79 38.75 75.12

 

67

 

 

 

 

 

 

APPENDR B (continued)
Ultramafics
A5G A521 ASH A81 A90

5102 43.62 40.67 43.49 41.61 42.30
A1203 5.12 3.94 5.49 4.58 5.00
Fe203* 1.83 1.94 1.88 1.97 1.78
FeO 12.18 12.96 12.50 13.11 11.90
MgO 30.79 36.76 31.62 35.45 33.19
CaO 4.91 2.47 3.59 2.00 4.19
NaZO .15 .10 .01 .07 .23
K20 .16 .19 .15 .17 .21
T102 1.02 .75 1.03 .81 .98
P205 .08 .07 .08 .07 .07
MnO .15 .16 .15 .15 .16
Norms

Q 0.00 0.00 0.00 0.00 0.00
C 0.00 0.00 0.00 .81 0.00
Or .95 1.12 .90 1.03 1.25
Ab 1.26 .85 .10 .59 1.97
An 12.82 9.75 14.47 9.46 11.96
Ne 0.00 0.00 0.00 0.00 0.00
Wo 4.60 .86 1.16 0.00 3.49
En 19.16 7.01 22.05 14.57 11.61
Fs 4.86 1.63 5.60 3.53 2.68
Mt 2.65 2.82 2.72 2.85 2.59
11 1.94 1.42 1.97 1.55 1.86
Ap .19 .16 .20 .17 .16
01 51.58 74. 39 50.85 65 .46 62.44
(Di) 8.78 1.64 2.21 0.00 6.65
(Hy) 19.84 7.86 26.59 18.09 11.13
Plag 91.06 91.97 99.32 94.13 85.84
Mg* 83.23 84.78 83.23 84.15 84.56
Fe203* = 0.15 FeO

Mg* = 100 x Mg/Mg + Fe

 

 

 

 

 

 

 

 

00.0 00.0 0.00 0.000 0.00 0.000 0.02 0.00 0.00 0.00 3 000.0
00.0 00; 0.00 0.000 040 0.000 0.02 0.00 0.00 0.00 0.0 _0<
00.0 00.0 .02
00.0 03 0.00 0.000 0.00 0.000 0.000 0.00 0.00 0.00 0.0 02
00.0 00; 0.00 0.000 0.00 0.0: 0.00_ 0.: 0.00 0.00 0.0 002
ohnnmwmuoz
2.0 00.2 0.00 0.0 0.000 0.00 0.000 0.0 00.0 0.: 0.0 02
00.0 00.: 0.02 0.0 0.000 0.00 0.000 0.0 00.0 0.2 0.0 02
ohxzmocmhu
00.. 00.0 0.02 0.000 0.000 0.20 0.000 0.00 0.00 0.2 0.0 EB
00.0 00.0 0.00 0.02 0.00 0.000 0.00 0.0: 0.00 0.00 0.0 002
00.0 00.0 0.000 0.000 0.000 0.000 0.0000 0.00 0.00 0.00 0._ 00000.
00.0 010 0.000 0.000 0.00 0.0: 0.000 0.02 0.0 0.0_ 0._ 008
00.0 00.0 0.00 0.000 0.00 0.000 0.000 0.00 0.: 0.00 0.0 98
004 00.0 0.000 0.000 0.02 0.02 0.00 0.00 0.0... 0.00 04 008
00.0 00.. 0.000 0.000 0.00 0.02 0.00 0.00 0.00 0.00 0.0 00000
00.0 03 0.00 0.000 0.0_ 0.000 0.00 0.02 0.00 0.00 0.0 000a
£83302
0.0 0: 00 > .0 .0 00 _z 5 R0 60 20500

 

 

<H<Q HZmEmAm mU<mh

U 502mm:

68

69

 

 

 

 

 

00.0 0.00 0.000 0.00 0.00 0.00 0.0000 0.00 00000
00.0 0.00 0.000 0.00 0.00 0.00 0.0000 0.00 00<
00.0 0.000 0.000 0.00 0.00 0.000 0.000 0.00 :2
00.0 0.00 0.00 0.00 0.00 0.00 0.0000 0.00 02
00.0 0.000 0.000 0.00 0.00 0.00 0.0000 0.00 02
8:000:82:
00.0 0.00 0.00 0.00 0.000 0.000 0.000 0.00 "00000
00.0 0.000 0.000 0.000 0.000 0.000 0.00 0.00 0000
00.0 0.00 0.000 0.00 0.000 0.000 0.00 0.00 <2
00.0 0.00 0.000 0.00 0.000 0.000 0.000 0.00 0000
00.0 0.00 0.000 0.00 0.000 0.00 0.00 0.00 00<
00.0 0.000 0.000 0.000 0.000 0.000 0.00 0.00 <0<
oEnmwmuoE
0» 00 > .0 .0 00 02 no 600500

 

 

0080000080 0 005200.090

7O

 

 

 

 

 

 

 

00.0 00.0 00.0 00.0 00.0 00.00 00.00 00.00 00.00 00.000 00.0 .0000.
00.0 00.0 00.0 00.0 00.0 00.00 00.00 00.00 00.00 00.00 00.0 000.
00.0 00.0 00.0 00.0 00.0 00.00 00.00 00.00 00.00 00.0 00.0 .0000
00.0 00.0 00.0 00.0 00.0 00.00 00.0 00.00 00.00 00.000 00.0 .0000
00.0 00.0 00.0 00.0 00.0 00.00 00.00 00.00 00.00 00.000 00.0 000<
oEnmJMEoE
00.0 00.00 00.0 00.0 00.00 00.000 00.000 00.0 00.0 00.0 00.00 0000
00.0 00.00 00.0 00.0 00.00 00.000 00.00 00.0 00.0 00.0 00.0 02
«ficaocmhu
00.0 00.0 00.0 00.0 00.00 00.00 00.00 00.00 00.00 00.00 00.0 000000
00.0 00.0 00.0 00.0 00.0 00.00 00.0 00.00 00.00 00.000 00.0 000000
00.0 00.0 00.0 00.0 00.0 00.00 00.00 00.00 00.00 00.00 00.0 000000
00.0 00.0 00.0 00.0 00.0 00.00 00.0 00.00 00.00 00.000 00.0 00000
00.0 00.0 00.0 00.0 00.0 00.00 00.00 00.00 00.00 00.000 00.0 0.0000
00.0 00.0 00.0 00.0 00.0 00.00 00.00 00.00 00.00 00.00 00.0 00000
00.0 00.0 00.0 00.0 00.0 00.00 00.00 00.00 00.00 00.000 00.0 00000
00.0 00.0 00.0 00.0 00.0 00.0 00.0 00.00 0.00 00.000 00.0 0000
00000000922
0.0 0» 00. am .00 60 3 oo 60 .6 00.0 000500

 

 

0060000080 0 00000200090

71

 

 

 

 

 

 

00.0 00.0 00.0 00.0 00.00 00.0 00.000 00.00 0.0000 00.0 00000
00.0 00.0 00.0 00.0 00.00 00.0 00.000 00.00 0.0000 00.0 000.
00.0 00.0 00.0 00.0 00.00 00.0 00.00 00.00 0.0000 00.0 0.000.
00.0 00.0 00.0 00.0 00.00 00.0 00.000 00.00 0.0000 00.0 00<
00.0 00.0 00.0 00.0 00.00 00.0 00.000 00.00 0.0000 00.0 0000
0039:0005
00.0 00.0 00.0 00.0 00.0 00.0 00.0 00.00 00.00 00.000 00.0 "00000
00.0 00.0 00.0 00.0 00.0 00.00 00.00 00.00 00.00 00.00 00.0 000000
00.0 00.0 00.0 00.0 00.0 00.00 00.0 00.00 00.00 00.000 00.0 <0<
00.0 00.0 00.0 00.0 00.0 00.00 00.00 00.00 00.00 00.000 00.0 0000
00.0 00.0 00.0 00.0 00.0 00.00 00.00 00.00 00.00 00.000 00.0 0000
00.0 00.0 00.0 00.0 00.0 00.00 00.00 00.00 00.00 00.00 00.0 <0<
ohnnmwmuoz
3 er 00.0 am 50 60 3 co 60 .u 0: 000.000

 

 

00000000000080 U 00520—0002

APPENDIX D

SAMPLE LOCATIONS AND DESCRIPTIONS

 

 

Hemlock Metabasalt

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

029B Pillow basalt, SE14 NEK, sec. 27, T43N R31W, 46° 6' 50", 88° 11' 27".

029C Massive basalt, location as above, uncertain relation to 0293.

029E Basalt, SW11 SW14, sec. 32, T43N R31W, 46° 4137.1", 88° 13' 12.1".

031D Pillow basalt (classified as tholeiitic andesite), gague contact with
sill, N15 SW16, sec. 20, T43N R31W, 46° 6' 31. 7", 88 12' 57.1".

D34G Amygdaloidal basalt, onorth of Michigamme Falls Dam, SE16 SE14,
sec. 7, T43N R31W, 46 7' 53. 6", 88° 13' 30.1".

D65A Basalt, SE16 SW14, sec. 25, T43N R32W, location uncertain.

D68A Basalt, west of road, NEK NW16, sec. 4, T43N R32W, 46° 9121.0",
88° 18' 58. 5.

D71A Basalt, an NWK, Sec. 30, T44N R32W, 46° 11' 17", 88° 21' 50".

Granophyre
A5B East and south of road, NW16 NWK, sec. 12, T43N R32W, 46° 8' 31.2",
15' 48.5".

A5C Location as above, about 10 yards from A5B, contact relation

covered.
Metagabbro
A532 MaPped as granophyre, swx NEK, sec. 11, T43N R32W, 46° 8' 18.3",
16' 11.8".

A51. As above.

A5P Mapped as granophyre, easct> of railroad tracks, W16 NWK, sec. 12,
T43N R32W, 46° 8' 28. 7", 88 15' 42.8".

A61 MaPped as granophyre, SEK swx, sec. 28, T4414 R32W, 46° 10' 32.6",

18' 57.1".

 

72

73

APPENDIX D (continued)

 

 

 

 

 

 

 

 

 

 

 

 

 

 

Metagabbro
A27F mapped as granophyre, NEK swn, sec. 7, T43N R31W, 46° 8' 7. 5",
13' 34. 8".
A2A Center of swx, sec. 32, T43N R31W, 46° 4' 38.6", 88° 12' 57.8".
A4B Basal sectiono 036 sill, im mgdiately west of road, NEK NWK, sec. 21,
T43N R31W, 46 5,0" 88 11'2 2.7"
A4C As above, about 100 yards from A4B.
A5A Associated with the granophyre, north of road, NWK NWK, sec. 12,
T43N R32W, 46° 8' 34. 1", 88° 15' 45. 7".
A29D SW14 SW14, sec. 32, T43N R31W, 46° 4137.1", 88° 13' 10.0".
A29F Layered outcro og (classifiedo as alkalic picrite), SWK NWK, sec. 16,
T43N R31W, 46 7' 28.0", 88° 11' 57.8".
Ultramafic Cumulate
A5C N16, sec. 22, T44N R32W, 46° 12' 7", 88° 17' 30", location uncertain.
A51-1 As above.
A9D As above.
A511 SK, sec. 15, T44N R32W, 46° 12' 27", 88° 17' 40".
A81 As above.

 

 

REFERENCES

Arth, J. G. and Hanson, G. N., 1975. Geochemistry and origin of the early
Precambrian crust of northeastern Minnesota. Geochim. Cosmochim.
Acta, v. 39, p. 325-362.

Banks, P. O. and Van Schmus, W. R., 1972. Chronology of Precambrian rocks of
iron and Dickinson Counties, Michigan, Part 11 (abstr.). 18th Ann. Inst.
Lake Superior Geol., Paper 23.

Baragar, W. R. A. and Scoates, R. F. J., 1981. The circum-Superior belt: a
Proterozoic plate margin? I_n Precambrian Plate Tectonics, A. Kroner
(ed.), Elsevier Publishing, p. 297-330.

Bayley, R. W., 1959a. A metamorphosed differentiated sill in northern Michigan.
Amer. Jour. Sci., v. 257, p. 408-430.

Bayley, R. W., 1959b. Geology of the Lake Mary quadrangle, Iron County,
Michigan. U.S. Geol. Surv. Bull. 1077, 112p.

Beccaluva, L., Ohnenstetter, D., Ohnenstetter, M. and Venturelli, G., 1977. The
trace element geochemistry of Corsican ophiolites. Contrib. Mineral.
PCtrOIo, V. 6", p. 11-310

Bhat, M. 1., Zaihuddin, S. M. and Rais, A., 1981. Panjal trap chemistry and the
birth of Tethys. Geol. Mag., v. 118, p. 367-375.

Burke, K. and Dewey, J. F., 1972. An outline of Precambrian plate development.
l_n_ Implications of Continental Drift to the Earth Sciences, D. H. Tarling
and S. K. Runcorn (eds.), Academic Press, p. 1035-1045.

Cambray, F. W., 1978. Plate tectonics as a model for the environment of
deposition and deformation of the early Proterozoic (Precambrian X) of
northern Michigan. Geol. Soc. Amer. Abstr. w/Programs, v. 10, p. 7.

Cannon, W. F., 1973. The Penokean Orogeny in northern Michigan. Geol. Assoc.
Canada, Spec. Pap. 12, p. 251-271.

Cannon, W. F. and Gair, J. E., 1970. A revision of stratigraphic nomenclature
for middle Precambrian rocks in northern Michigan. Geol. Soc. Amer.
BUllo, V. 81, p. 28l‘3-28460

Cannon, W. F. and Klasner, J. S., 1975. Stratigraphic relationships within the
Baraga Group of Precambrian age, central Upper Peninsula, Michigan. U.S.
C1601. SUI’V- 3001‘. R850, V. 3, p. (87-510

Cannon, W. F. and Klasner, J. S., 1976. Geologic map and geophysical

interpretation of the Witch Lake quadrangle, Marquette, Iron and Baraga
Counties, Michigan. Misc. Geol. Invest. Map 1-987.

74

75

Condie, K. C. and Baragar, W. R. A., 1974. Rare-earth element distributions in

volcanic rocks from Archean Greenstone Belts. Contrib. Mineral. Petrol.,
v. 45, p. 237-246.

Condie, K. C., Viljoen, M. J. and Kable, E. J. D., 1977. Effects of alteration on
element distributions in Archean tholeiites from the Barberton Greenstone
Belt, South Africa. Contrib. Mineral. Petrol., v. 64, p. 75-89.

Cudzilo, T. F., 1978. Geochemistry of early Proterozoic igneous rocks,
northeastern Wisconsin and upper Michigan. Unpub. Ph.D. Dissertation,
Univ. of Kansas, 202p.

Dutton, C. E., 1971. Geology of the Florence area, Wisconsin and Michigan.
US. Geol. Surv. Prof. Pap. 633, 54p.

Dutton, C. E. and Linebaugh, R. E., 1967. Map showing Precambrian geology of
the Menominee iron-bearing district and vicinity, Michigan and Wisconsin.
Misc. Geol. Invest. Map 1-466.

Ferrara, G., Innocenti, F., Ricci, C. A. and Serri, G., 1976. Ocean-floor affinity
of basalts from north Apennine ophiolites: geochemical evidence. Chem.
Geol., v.17, p. 101-111.

Floyd, P. A. and Winchester, J. A., 1975. Magma type and tectonic setting
discrimination using immobile elements. Earth Planet. Sci. Lett., v. 27,
p. 211-218.

Frey, F. A., Haskin, M. A., Poetz, J. A. and Haskin, L. A., 1968. Rare earth
abundances in some basic rocks. J. Geophys. Res., v. 73, p. 6085-6098.

Frey, F. A., Haskin, L. A. and Haskin, M. A., 1971. Rare-earth abundances in
some ultramafic rocks. J. Geophys. Res., v. 76, p. 2057-2070.

Frey, F. A., Bryan, W. B. and Thompson, G., 1974. Atlantic Ocean floor:
geochemistry and petrology of basalts from legs 2 and 3 of the deep-sea
drilling project. J. Geophys. Res., v. 79, p. 5507-5527.

Frey, F. A., Green, D. H. and Roy, S. D., 1978. Integrated models of basalt
petrogenesis: a study of quartz tholeiites to olivine melilitites from

southeastern Australia utilizing geochemical and experimental petrological
data. J. Petrol., v. 19, p. 463-513.

Gair, J. E. and Wier, K. L., 1956. Geology of the Kiernan quadrangle, Iron
County, Michigan. U.S. Geol. Surv. Bull. 1044, 88p.

Gibb, R. A., Thomas, M. D. amd Mukhopadhyay, M., 1980. Proterozoic sutures in
canadao GOOSCI- cane, V0 7, p0 ll‘9-154o

Gill, R. C. O., 1979. Comparative petrogenesis of Archean and modern low-K
tholeiites, a critical review of some geochemical aspects. In Origin and
Distribution of the Elements, L. H. Ahrens (ed.), Pergamon Press,
p. 431-447.

76

Gill, R. C. O. and Bridgwater, D., 1976. The Ameralik dykes of West Greenland,
the earliest known basaltic rocks intruding stable continental crust. Earth
Planet. Sci. Lett., v. 29, p. 276-282.

Gordon, G. E., Randle, K., Goles, G. G., Corliss, J. B., Beeson, M. H. and
Oxley, S. 5., 1968. Instrumental activation analysis of standard rocks with

high-resolution gamma-ray detectors. Geochim. Cosmochim. Acta, v. 32,
p. 369-396.

Hart, 5. R., Brooks, C., Krogh, T. E., Davis, G. L. and Nava, D., 1970. Ancient
and modern volcanic rocks: a trace element model. Earth Planet. Sci.
Lett., V0 IO, p. 17-280

Haskin, L. A. and Haskin, M. A., 1968. Rare earth elements in the Skaergaard
intrusion. Geochim. Cosmochim. Acta, v. 32, p. 433-447.

Hellman, P. L. and Henderson, P., 1977. Are rare earth elements mobile during
spilitisation? Nature, v. 267, p. 38-40.

Hellman, P. L., Smith, R. E. and Henderson, P., 1977. Rare earth element

investigation of the Cliefden Outcrop, N.S.W., Australia. Contrib. Mineral.
PetrOIO, V. 65, p. 155-164.

Hellman, P. L., Smith, R. E. and Henderson, P., 1979. The mobility of the rare
earth elements: evidence and implications from selected terrains affected
by burial metamorphism. Contrib. Mineral. Petrol., v. 71, p. 23-44.

Helmke, P. A. and. Haskin, L. A., 1973. Rare-earth elements, Co, Sc, and Hf in
the Steens Mountain basalts. Geochim. Cosmochim. Acta, v. 37,
p. 1513-1529.

Herrman, A. G., Potts, M. J. and Knake, D., 1974. Geochemistry of the rare
earth elements from the oceanic and continental crust. Contrib. Mineral.
PCtl’Ole, V0 4", p0 1‘16.

Hoffman, P., 1973. Evolution of an early Proterozoic continental margin: the
Coronation geosyncline and associated aulacogens of the northwestern
Canadian shield. Phil. Trans. R. Soc. Lond., v. 273A, p. 547-581.

Humphris, S. E. and Thompson, G., 1978a. Hydrothermal alteration of oceanic
basalts by seawater. Geochim. Cosmochim. Acta, v. 42, p. 107-125.

Humphris, S. E. and Thompson, G., 1978b. Trace element mobility during
hydrothermal alteration of oceanic basalts. Geochim. Cosmochim. Acta,
V0 (*2, p. 127-136.

Humphris, S. E., Morrison, M. A. and Thompson, R. N., 1978. Influence of rock
crystallization history upon subsequent lanthanide mobility during
hydrothermal alteration of basalts. Chem. Geol., v. 23, p. 125-137.

Irvine, T. N. and Baragar, W. R. A., 1971. A guide to the chemical classification
of the common volcanic rocks. Can. Jour. Earth Sci., v. 8, p. 523-548.

77

of Archean volcanic rocks. Geochim. Cosmochim. Acta, v. 38, p. 611-627.

James, H. L., Clark, L. D., Lamey, C. A. and Pettijohn, F. J., 1961. Geology of
central Dickinson County, Michigan. U.S. Geol. Surv. Prof. Paper 310,
176p.

James, H. L., Dutton, C. E., Pettijohn, F. J. and Wier, K. L., 1968. Geology and
ore deposits of the Iron River - Crystal Falls District, Iron County,
Michigan. U.S. Geol. Surv. Prof. Paper 570, 134p.

Kay, R. W. and Hubbard, N. J., 1978. Trace elements in ocean ridge basalts.
Earth Planet. SCio Lett., V. 38, p. 95-116.

Kay, R. W. and Senechal, R. G., 1976. Rare earth geochemistry of the Troodos
ophiolite complex. J. Geophys. Res., v. 81, p. 964-970.

Kuo, H. Y. and Crockett, J. H., 1979. Rare earth elements in the Sudbury nickel
irruptive: comparison with layered gabbros and implications for nickel
irruptive petrogenesis. Econ. Geol., v. 74, p. 590-605.

Larue, D. K., 1981. The early Proterozoic pre-iron-formation Menominee Group
siliciclastic sediments of the southern Lake Superior region: evidence for
sedimentation in platform and basinal settings. J. Sed. Petrol., v. 51,
p. 397-414.

Larue, D. K., 1981. Sedimentary, structural, and chemical evolution of a
purported continental margin of early Proterozoic age, Lake Superior
region. N.S.F. grant number EAR-81-08514.

LeBlanc, M., 1981. The late Proterozoic ophiolites of Bou Azzer (Morocco):
evidence for Pan-African plate tectonics. E Precambrian Plate Tectonics,
A. Kroner (ed.), Elsevier Publishing, p. 435-447.

Leeman, W. P., 1977. Comparison of Rb/Sr, U/Pb, and rare earth element
characteristics of sub—continental and sub-oceanic mantle regions. AGU
Chapman Conf. on partial melting in the upper mantle, Oregon Dept. of
Geol. and Min. Indus. Bull. 96, p. 149-168.

Ludden, J. N. and Thompson, G., 1978. Behavior of rare earth elements during
submarine weathering of tholeiitic basalt. Nature, v. 274, p. 147-149.

Ludden, J. N. and Thompson, G., 1979. An evaluation of the behavior of the rare
earth elements during the weathering of sea-floor basalt. Earth Planet.
Sci. Lett., v. 43, p. 85-92.

Masuda, A. and Nagasawa, S., 1975. Rocks with negative cerium anomalies,
dredged from Shatsky Rise. Geochem. Jour., v. 9, p. 227-233.

Menzies, M. and Moores, E., 1976. The nature and origin of the oceanic crust:
rare earth and morphological evidence from ophiolites. Amer. Geophys.
union TranSO’ V. 57’ p0 (PCB-404.

78

Menzies, M., Blanchard, D. and Jacobs, J., 1977. Rare earth and trace element
geochemistry of metabasalts from the Point Sal ophiolite, California.
Earth Planet. Sci. Lett., v. 37, p. 203-215.

Menzies, M., Seyfried, W. and Blanchard, D., 1979. Experimental evidence of
rare earth element immobility in greenstones. Nature, v. 282, p. 398-399.

Menzies, M., Blanchard, D. and Xenophontos, C., 1980. Genesis of the
Smartville arc-ophiolite, Sierra Nevada foothills, California. Amer. Jour.
SCio, V. 280A, p0 329-34“.

Montigny, R., Bougault, H., Bottinga, Y. and Allegre, C. J., 1973. Trace element
geochemistry and genesis of the Pindos ophiolite suite. Geochim.
COSTDOChImo ACta, V. 37, p. 2135-2147.

Morey, G. B., 1978. Metamorphism in the Lake Superior region, USA, and its
relation to crustal evolution. Geol. Surv. Can. Paper 78-10, p. 283-314.

Morey, G. B. and Sims, P. K., 1976. Boundary between two Precambrian W
terranes in Minnesota and its geologic significance. Geol. Soc. Amer. Bull.
87, P. 141-152.

Morris, W. J., 1977. Geochemistry and origin of the Yellow Dog Plains
peridotite, Marquette County, northern Michigan. Unpub. M.S. Thesis,
Michigan State Univ., 75p.

Morrison, M. A., 1978. The use of "immobile" trace elements to distinguish the
paleotectonic affinities of metabasalts: applications to the Paleocene
basalts of Mull and Skye, northwest Scotland. Earth Planet. Sci. Lett.,
v. 39, p. 407-416.

Muecke, G. K., Pride, C. and Sarkar, P., 1979. Rare-earth element geochemistry
of regional metamorphic rocks. I_n_ Origin and Distribution of the Elements,
L. H. Ahrens (ed.), Pergamon Press, p. 449-464.

Nesbitt, R. W. and Sun, S.-S., 1976. Geochemistry of Archean spinifex-textured
peridotites and magnesian and low-Mg tholeiites. Earth Planet. Sci. Lett.,
V0 31, p. 433-453.

Noiret, G., Montigny, R. and Allegre, C. J., 1981. Is the Vourinos Complex an
island arc ophiolite? Earth Planet. Sci. Lett., v. 56, p. 375-386.

O'Nions, R. K. and Pankhurst, R. J., 1978. Early Archean rocks and geochemical
evolution of the earth's crust. Earth Planet. Sci. Lett., v. 38, p. 211-236.

Pallister, J. S. and Knight, R. J., 1981. Rare earth element geochemistry of the
Samail ophiolite near Ibra, Oman. Jour. Geophys. Res., v. 86,
p0 2673-26970

Paster, T. P., Schauwecker, D. S. and Haskin, L. A., 1974. The behavior of some
trace elements during solidification of the Skaergaard layered series.
Geochim. Cosmochim. Acta, v. 38, p. 1549-1577.

79

Pearce, J. A. and Cam, J. R., 1971. Ophiolite origin investigated by
discriminant analysis using T1, Zr, and Y. Earth Planet. Sci. Lett., v. 12,
p0 339-3l‘90

Pearce, J. A. and Cann, J. R., 1973. Tectonic setting of basic volcanic rocks
determined using trace element analysis. Earth Planet. Sci. Lett., v. 19,
P0 290-3000

Pearce, J. A. and Norry, M. J., 1979. Petrogenetic implications of Ti, Zr, Y and
Nb variations in volcanic rocks. Contrib. Mineral. Petrol., v. 69, p. 33-47.

Pearce, T. H., Gorman, B. E. and Birkett, T. C., 1975. The TiOZ-KZO-P205
diagram: a method of discriminating between oceanic and non-oceanic
basalts. Earth Planet. Sci. Lett., v. 24, p. 419-426.

Pearce, T. H., Gorman, B. E. and Birkett, T. C., 1977. The relationship between
major element chemistry and tectonic environment of basic and
intermediate volcanic rocks. Earth Planet. Sci. Lett., v. 36, p. 121-132.

Phelps, D., 1980. The Sparta ophiolite complex, northeast Oregon: a plutonic
equivalent to low K20 island-arc volcanism. Amer. Jour. Sci., v. 280A,
p. 345-358.

Philpotts, J. A. and Schnetzler, C. C., 1968. Europium anomalies and the genesis
Of basalt. Chemo COOL, V. 3, P. 5.13.

Philpotts, J. A., Schnetzler, C. C. and Hart, 5. R., 1969. Submarine basalts:
some K, Rb, Sr, Ba, rare-earth, H O and CO data bearing on their
alteration, modification by plagioclase and possib e source material. Earth
Planet. Sci. Lett., v. 7, p. 293-299.

Potts, M. J. and Condie, K. C., 1971. Rare earth element distributions in a

proto-stratiform ultramafic intrusion. Contrib. Mineral. Petrol., v. 33,
P. Zle'258o

Robertson, A. H. F. and Fleet, A. J., 1976. The origins of rare earths in
metalliferous sediments of the Troodos Massif, Cyprus. Earth Planet. Sci.
Lett., v. 28, p. 385-394.

Schilling, J. G., 1971. Sea-floor evolution: rare-earth evidence. Phil. Trans. R.
SOC. 1.0qu V0 268A, p. 663-706.

Serri, G., 1981. Petrochemistry of ophiolite gabbroic complexes: a key for the
classification of ophiolites into Lo-Ti and Hi-Ti types. Earth Planet. Sci.
Lett., V0 51, p. 203-212.

Shervais, J. W., 1982. Ti-V plots and the petrogenesis of modern and ophiolitic
lavas. Earth Planet. Sci. Lett., v. 59, p. 101-118.

Sims, P. K., 1976. Precambrian tectonics and mineral deposits, Lake Superior
region. ECO“. 58010, V0 71, P0 1092-11270

8O

Sims, P. K., Card, K. D., Morey, G. B. and Peterman, 2. E., 1980. The Great
Lakes tectonic zone - a major crustal structure in North America. Geol.
SOC- Amero BUllo, V. 91, p. 690-698.

Smewing, J. D. and Potts, P. J., 1976. Rare-earth abundances in basalts and
metabasalts from the Troodos Massif, Cyprus. Contrib. Mineral. Petrol.,
V0 57, po 245-2580

Smith, R. E. and Smith, S. E., 1976. Comments on the use of Ti, Zr, Y, Sr, K, P,
and Nb in classification of basaltic magmas. Earth Planet. Sci. Lett.,
V. 32, p. 114-120.

Stolper, E., 1980. A phase diagram for mid-ocean ridge basalts: preliminary
results and implications for petrogenesis. Contrib. Mineral. Petrol., v. 74,
p. 13-27.

Strong, D. F. and Dostal, J., 1978. Proterozoic oceanic crust in the Avalon zone
of Newfoundland. Geol. Soc. Amer. Abstr. w/Programs, v. 10, p. 500.

Strong, D. F. and Dostal, J., 1980. Dynamic melting of Proterozoic upper
mantle: evidence from rare earth elements in oceanic crust of eastern
Newfoundland. Contrib. Mineral. Petrol., v. 72, p. 165-173.

Strong, D. F., O'Brien, 5. J., Taylor, S. W., Strong, P. G. and Wilton, D. H., 1978.
Aborted Proterozoic rifting in eastern Newfoundland. Can. Jour. Earth
Sci., v. 15, p. 117-131.

Suen, C. J., Frey, F. A. and Malpas, J., 1979. Bay of Islands ophiolite suite,
Newfoundland: petrologic and geochemical characteristics with emphasis

on rare earth element geochemistry. Earth Planet. Sci. Lett., v. 45,
p. 337-348.

Sun, S.-S. and Nesbitt, R. W., 1977. Chemical heterogeneity of the Archean

mantle, composition of the earth and mantle evolution. Earth Planet. Sci.
Lett., V. 35, P0 a29-l‘48.

Sun, S.-S. and Nesbitt, R. W., 1978. Petrogenesis of Archean ultrabasic and basic

volcanics: evidence from rare earth elements. Contrib. Mineral. Petrol.,
V0 65, p. 301-325.

Sun, S.-S., Nesbitt, R. W. and Sharaskin, A. Y., 1979. Geochemical
characteristics of mid-ocean ridge basalts. Earth Planet. Sci. Lett., v. 44,
p0 119‘1380

Upadhyay, H. D. and Neale, E. R. W., 1979. On the tectonic regimes of ophiolite
genesis. Earth Planet. Sci. Lett., v. 43, p. 93-102.

Van Schmus, W. R., 1974. Chronology of Precambrian events in Wisconsin.
Amer. Geophys. Union Trans., v. 55, p. 465.

Van Schmus, W. R., 1976. Early and middle Proterozoic history of the Great
Lakes area, North America. Phil. Trans. R. Soc. Lond., v. 280A,
p. 605-628.

81

Van Schmus, W. R. and Anderson, J. L., 1977. Gneiss and migmatite of Archean
age in the Precambrian basement of central Wisconsin. Geology, v. 5,
p0 #5-l‘8.

Van Schmus, W. R., Thurman, E. M. and Peterman, Z. E., 1975. Geology and
Rb-Sr chronology of middle Precambrian rocks in eastern and central
Wisconsin. Geol. Soc. Amer. Bull., v. 86, p. 1255-1265.

Venturelli, G., Capedri, 5., Thorpe, R. S. and Potts, P. J., 1979. Rare-earth and
other element distribution in some ophiolitic metabasalts of Corsica,
western Mediterranean. Chem. Geol., v. 24, p. 339-353.

Walker, D., Shibata, T. and DeLong, S. E., 1979. Abyssal tholeiites from the
Oceanographer fracture zone 11 - phase equilibria and mixing. Contrib.
Mineral. Petrol., v. 70, p. 111-125.

Weill, D. F. and Drake, M. J., 1973. Europium anomaly in plagioclase feldspar:
experimental results and semi-quantitative model. Science, v. 180,
p0 1059-10600

Wier, K. L., 1967. Geology of the Kelso Junction quadrangle, Iron County,
Michigan. U.S. Geol. Surv. Bull. 1226, 47p.

Wilband, J. T. and Wasuwanich, P., 1980. Models of basalt petrogenesis: lower
Keweenawan diabase dikes and middle Keweenawan Portage Lake Lavas,
upper Michigan. Contrib. Mineral. Petrol., v. 75, p. 395-406.

Winchester, J. A. and Floyd, P. A., 1977. Geochemical discrimination of
different magma series and their differentiation products using immobile
elements. Chemo GeOIO, V0 20, p. 325'3g30

Windley, B. F., 1973. Crustal development in the Precambrian. Phil. Trans. R.
Soc. Lond, v. 273A, p. 321-341.

Wood, D. A., 1978. Major and trace element variations in the Tertiary Lavas of
eastern Iceland and their significance with respect to the Icelandic
geochemical anomaly. J. Petrol., v. 19, p. 393-436.

Wood, D. A., 1980. The application of a Th-Hf-Ta diagram to problems of
tectonomagmatic classification and To establishing the nature of crustal

contamination of basaltic lavas of the British Tertiary volcanic province.
Earth Planet. Sci. Lett., v. 50, p. 11-30.

Wood, D. A., Gibson, 1. L. and Thompson, R. N., 1976. Elemental mobility during
zeolite facies metamorphism of the Tertiary basalts of eastern Iceland.
Contrib. Mineral. Petrol., v. 55, p. 241-254.

Wood, D. A., Joron, J.-L. and Treuil, M., 1979. A reappraisal of the use of trace
elements to classify and discriminate between magma series erupted in
different tectonic settings. Earth Planet. Sci. Lett., v. 45, p. 326-336.