U.S. Department of the Interior

U.S. Geological Survey

Professional Paper 1730

The Gogebic Iron Range—A Sample of the 

Northern Margin of the Penokean Fold and 

Thrust Belt
Cover: Ironwood Iron-Formation from Vicar Mine near Wakefield, Mich. This section 

shows red granular jasper interbedded with specular hematite and magnetite. Scale is 

approximately 1:1.
The Gogebic Iron Range—A Sample of 

the Northern Margin of the Penokean 

Fold and Thrust Belt

By William F. Cannon, Gene L. LaBerge, John S. Klasner, and Klaus J. Schulz

Professional Paper 1730

U.S. Department of the Interior

U.S. Geological Survey
U.S. Department of the Interior

DIRK KEMPTHORNE, Secretary

U.S. Geological Survey

Mark D. Myers, Director

U.S. Geological Survey, Reston, Virginia: 2008

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Suggested citation:

Cannon, W.F., LaBerge, G.L., Klasner, J.S., and Schulz, K.J. 2007, The Gogebic iron range—a sample of the northern 

margin of the Penokean fold and thrust belt: U.S. Geological Survey Professional Paper 1730, 44 p. 

Library of Congress Cataloging-in-Publication Data

The Gogebic iron range : a sample of the northern margin of the Penokean fold 

and thrust belt / by William F. Cannon . . . [et al.].

 p. cm. -- (Professional paper ; 1730)

 Includes bibliographical references.

 ISBN-13 : 978-1-4113-1917-2

1. Geology, Stratigraphic--Proterozoic. 2. Geology--Gogebic Range 

Region (Mich. and Wis.) 3. Folds (Geology)--Gogebic Range Region (Mich. 

and Wis.) 4. Orogenic belts--Gogebic Range Region (Mich. and Wis.) 

I. Cannon, William F.

 QE653.5.G65 2007

 557.74--dc22 2007029280
iii

Contents

Abstract...........................................................................................................................................................1

Introduction.....................................................................................................................................................2

General Geology.............................................................................................................................................4

Neoarchean Rocks...............................................................................................................................4

Paleoproterozoic Rocks.......................................................................................................................4

Chocolay Group............................................................................................................................6

Sunday Quartzite.................................................................................................................6

Bad River Dolomite.............................................................................................................6

Menominee Group.......................................................................................................................8

Palms Formation .................................................................................................................8

Ironwood Iron-Formation ..................................................................................................8

Relation between the Ironwood Iron-Formation and 

the Emperor Volcanic Complex...........................................................................9

Stratigraphic Relations in the Menominee Group and 

Emperor Volcanic Complex................................................................................10

Baraga Group .............................................................................................................................11

Copps Formation ...............................................................................................................12

Tyler Formation..................................................................................................................12

Pabst Member...................................................................................................................12

Age and Correlation of the Marquette Range Supergroup ................................................13

Stratigraphic Correlations within the Marquette Range Supergroup .....................13

Absolute Age of the Marquette Range Supergroup ...................................................14

Paleoproterozoic Intrusive Rocks...........................................................................................15

Mesoproterozoic Rocks.....................................................................................................................19

Structural Geology.......................................................................................................................................19

Mesoproterozoic Structures.............................................................................................................19

Atkins Lake-Marenisco Fault...................................................................................................21

Paleoproterozoic Structures.............................................................................................................22

Central Gogebic Iron Range.....................................................................................................23

Eastern Gogebic Iron Range....................................................................................................24

Extensional Phase ............................................................................................................24

Contractional Phase.........................................................................................................25

Summary of Eastern Gogebic Iron Range Structure ..................................................27

Western Gogebic Iron Range ..................................................................................................27

Economic Geology.......................................................................................................................................31

Natural Iron Ores ................................................................................................................................31

Taconite ................................................................................................................................................33

Synopsis ........................................................................................................................................................34

Sedimentation of the Marquette Range Supergroup ...................................................................34

Environmental Setting of Deposition ......................................................................................34

Chocolay Group.................................................................................................................34

Menominee Group............................................................................................................34

Baraga Group ....................................................................................................................35
iv

Tectonic Setting of Deposition.................................................................................................35

Chocolay Group.................................................................................................................35

Menominee Group............................................................................................................35

Baraga Group ....................................................................................................................36

Volcanism...........................................................................................................................36

Penokean Deformation ......................................................................................................................38

Acknowledgments.......................................................................................................................................40

References Cited..........................................................................................................................................40

Plates

(In pocket)

1. Geologic maps of the Gogebic iron range.

Figures

1. Generalized geologic map of the upper Great Lakes region, U.S. and Canada.................3

2. Schematic stratigraphic section along approximately 150 km of the Gogebic 

iron range.......................................................................................................................................5

3. Geologic map of part of the eastern Gogebic Range .............................................................7

3–4. Photographs showing

4. Stromatolites in Bad River Dolomite.................................................................................8

5. Examples of typical rocks types of the Emperor Volcanic Complex. .......................10

6. Schematic longitudinal section of part of the central Gogebic iron range ......................13

7. Geochemical discrimination diagrams showing the subalkaline tholeiitic 

character of the mafic rocks from the Emperor Volcanic Complex and diabase near 

Atkins Lake...................................................................................................................................14

8. Chondrite-normalized rare earth element plot and primitive-mantlenormalized multi-element plot for mafic rocks from the Emperor Volcanic

Complex and diabase near Atkins Lake..................................................................................18

9. Block diagram illustrating the geometric relationship between Neoarchean, 

Paleoproterozoic and Mesoproterozoic rocks in the central part of the Gogebic 

iron range.....................................................................................................................................22

10. Photomicrographs of rocks from the Atkins Lake-Marenisco fault zone on the 

north side of the Marengo River ..............................................................................................23

11. Schematic block diagram illustrating the geometry of fault sets in the central

Gogebic iron range.....................................................................................................................25

12. Diagram illustrating the restoration of the Paleoproterozoic geometry of the 

Presque Isle trough constructed by compensating for the Mesoproterozoic 

northward rotation......................................................................................................................27

13. Small scale recumbent fold in Ironwood Iron-Formation in NE1⁄4, sec. 24, T. 44 N., 

R. 4 W. ..........................................................................................................................................30

14. Schematic block diagram showing the relation of natural ores to the stratigraphy 

and structure of the central Gogebic iron range...................................................................32
v

15. Primitive-mantle-normalized multi-element plot showing the compositional 

similarity of mafic rocks from the Emperor Volcanic Complex and diabase near 

Atkins Lake with basalts of the Hemlock Formation and Badwater Greenstone ............37

16. Tectonic discrimination diagrams showing the within-plate affinity of the mafic 

rocks of the Emperor Volcanic Complex and diabase from near Atkins Lake .................38

17. Primitive-mantle-normalized multi-element plot showing the compositional 

similarity of rift-related Late Silurian to Early Devonian basalts from the northern 

Appalachians and the mafic rocks from the Emperor Volcanic Complex and 

diabase near Atkins Lake ..........................................................................................................39

Tables

1. Whole-rock chemical analyses of Paleoproterozoic mafic rocks from the 

Emperor Volcanic Complex and diabase near Atkins Lake.................................................16

2. Sequence of tectonic events in the Gogebic iron range. ....................................................20

The Gogebic Iron Range—A Sample of the Northern 

Margin of the Penokean Fold and Thrust Belt

By William F. Cannon,1 Gene L. LaBerge,2 John S. Klasner,3 and Klaus J. Schulz1

1

 U.S. Geological Survey, Reston, VA, 20192

2

 University of Wisconsin-Oshkosh and U.S. Geological Survey

3

 Western Illinois University and U.S. Geological Survey

Abstract

The Gogebic iron range is an elongate belt of Paleoproterozoic strata extending from the west shore of Lake Gogebic 

in the upper peninsula of Michigan for about 125 km westward into northern Wisconsin. It is one of six major informally 

named iron ranges in the Lake Superior region and produced 

about 325 million tons of direct-shipping ore between 1887 

and 1967. A significant resource of concentrating-grade ore 

remains in the western and eastern parts of the range.

The iron range forms a broad, gently southward-opening 

arc where the central part of the range exposes rocks that were 

deposited somewhat north of the eastern and western parts. 

A fundamental boundary marking both the tectonic setting 

of deposition and the later deformation within the Penokean 

orogen lies fortuitously in an east-west direction along the 

range so that the central part of the range preserves sediments deposited north of that boundary, whereas the eastern 

and western parts of the range were deposited south of the 

boundary. Thus, the central part of the range provides a record 

of sedimentation and very mild deformation in a part of the 

Penokean orogen farthest from the interior of the orogen to the 

south. The eastern and western parts of the range, in contrast, 

exhibit a depositional and deformational style typical of parts 

closer to the interior of the orogen. A second fortuitous feature 

of the iron range is that the entire area was tilted from 40° to 

90° northward by Mesoproterozoic deformation so that the 

map view offers an oblique cross section of the Paleoproterozoic sedimentary sequence and structures. Together, these 

features make the Gogebic iron range a unique area in which 

to observe (1) the lateral transition from deposition on a stable 

platform to deposition in a tectonically and volcanically active 

region, and (2) the transition from essentially undeformed 

Paleoproterozoic strata to their folded and faulted equivalents.

Paleoproterozoic strata in the Gogebic iron range are part 

of the Marquette Range Supergroup. They were deposited 

unconformably on Neoarchean rocks consisting of a diverse 

volcanic suite (the Ramsay Formation) which was intruded by 

granitic rocks of the Puritan Quartz Monzonite. The Marquette 

Range Supergroup in this region consists of a basal sequence 

of orthoquartzite (Sunday Quartzite) and dolomite (Bad River 

Dolomite), both of which are part of the Chocolay Group. 

The group is preserved only in the eastern and western parts 

of the range but was probably present throughout before the 

erosion interval that separated it from the overlying Menominee Group. The Menominee Group consists of basal clastic 

rocks (Palms Formation) that grade upward into the Ironwood 

Iron-Formation, which is the principal iron-bearing unit of the 

range. The Ironwood interfingers with the Emperor Volcanic 

Complex in the eastern part of the range and with volcanic 

rocks and gabbro in the western part of the range. The Ironwood is overlain unconformably by the Tyler Formation in 

the central and western parts of the range and by the Tyler’s 

equivalent, the Copps Formation, in the eastern part of the 

range.

Strata in the central part of the iron range are entirely 

sedimentary. Deposition occurred in a relatively stable tectonic 

setting, at least until the deposition of the Tyler Formation. 

The Tyler consists largely of turbidites deposited in a foreland basin in advance of accreting volcanic arcs to the south. 

Penokean deformation in the central part of the range was very 

minor; the evidence of deformation consists of steep faults 

with small offsets and a few bedding-parallel faults that also 

have small offsets and that are recognized only in mine workings. In both the eastern and western parts of the iron range, 

abrupt facies changes mark a passage into a more tectonically 

and volcanically active belt. These relationships are especially 

well displayed in the east where a graben, the Presque Isle 

trough, began to subside during deposition of the Ironwood 

Iron-Formation. The thickness of the Ironwood increases into 

the graben and its internal stratigraphy also changes. The most 

prominent changes in the graben are the presence of a thick 

volcanic unit, the Emperor Volcanic Complex of the Menominee Group, and comagmatic gabbro sills that interfinger with 

the Ironwood. In the western part of the range, volcanic rocks 

and comagmatic gabbro sills are also present in the Ironwood, 

but a graben that is equivalent to the Presque Isle trough is not 

evident.
2 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

Penokean structures are well developed in both the 

eastern and western parts of the iron range. They consist of 

folds ranging from outcrop to regional scale and thrust faults 

which, in places, either repeated the section or detached it 

from Neoarchean basement. The sharp transition from the 

little-deformed central part of the range to the more intensely 

deformed eastern and western parts coincides closely with the 

earlier developed transition from the stable sedimentary setting 

in the central part to the tectonically active sedimentation in 

the east and west parts. The extensional structures that formed 

during sedimentation may have helped to control the extent of 

later Penokean compressional structures.

Introduction

The Gogebic iron range is one of several geologically 

classic and economically significant iron ranges of the Lake 

Superior region (fig. 1). The Gogebic iron range is an arcuate 

east-west-trending belt in northern Michigan and Wisconsin 

that spans the northern margin of the Penokean fold-and-thrust 

belt and thus includes rocks deformed by the 1880–1830 Ma 

Penokean orogeny as well as rocks that remained largely undeformed. Paleoproterozoic metasedimentary and metavolcanic 

rocks, known collectively as the Marquette Range Supergroup, 

make up the iron range, but bounding Neoarchean basement 

rocks to the south and Mesoproterozoic rocks of the Midcontinent Rift on the north are also discussed. 

The Gogebic iron range has been studied and described 

in several generations of investigations, largely because of its 

former importance as a major iron-mining district. Studies 

of the entire range include Irving and Van Hise (1892), Allen 

and Barrett (1915) and Aldrich (1929). Numerous more local 

studies followed these early investigations. The intent of this 

paper is to summarize previous studies and add new information based on studies by the authors between 1990 and 1996, 

particularly in the eastern and western parts of the range; and 

to present a new stratigraphic and structural synthesis for the 

Paleoproterozoic strata and the Penokean foreland tectonic 

events that deformed them. Special emphasis is placed on the 

stratigraphic and structural relationships across the Penokean 

deformational boundary. The manuscript was prepared in 2003 

and in the intervening period until publication in 2008 several 

advances in understanding the regional setting of the Gogebic 

iron range have been made, which would modify some of the 

interpretive parts of this paper. But details of relationships 

within the range are not substantially changed since 2003.

The Paleoproterozoic strata of the Marquette Range 

Supergroup are at least partly correlative with the Huronian 

Supergroup of Ontario and the Animikie Group and Mille 

Lacs Group of Minnesota. Together, these sequences constitute an assemblage of sedimentary and lesser volcanic rocks 

deposited on Neoarchean basement rocks along the south-facing continental margin of the Superior Province (fig. 1). They 

record the transition from stable cratonic sedimentation in the 

oldest units, through extensional continental margin sedimentation, to deposition in a foreland basin during accretion of 

volcanic arcs in the youngest units (see Morey, 1996, for a 

summary of Paleoproterozoic stratigraphy of the region). All 

of these rocks now lie within or immediately north of the foldand-thrust belt of the Penokean orogen, which formed north of 

a suture marking the boundary of continental-margin rocks on 

the north with exotic volcanic terranes on the south.

In this paper, the Gogebic iron range is divided into three 

segments and each is discussed separately. The eastern segment extends from the easternmost exposures of the Copps 

Formation along the western shore of Lake Gogebic, westward 

to the Sunday Lake fault near Wakefield, Mich. (pl. 1B). The 

central segment, which includes all of the productive ironore district, continues from the Sunday Lake fault westward 

to Tyler Forks River near Upson, Wis. The western segment 

extends to the westernmost exposures of Paleoproterozoic 

strata where they are truncated by the Atkins Lake-Marenisco 

fault (pl. 1A, B). This report summarizes the geology of the 

central part of the iron range based mostly on previous studies, 

and presents new data and interpretations for the eastern and 

western parts of the range by combining previously published 

information with new field investigations conducted from 

1990 to 1996. This report also documents that substantial 

changes in both sedimentation and structure occurred over 

distances of only a few kilometers along strike. In particular, 

sedimentation of the major iron-bearing section, including the 

Ironwood Iron-Formation, is shown to span a tectonic boundary and to pass laterally through an interfingering relationship 

into volcanic rocks of the Emperor Volcanic Complex in the 

east.

Because the Gogebic iron range lies obliquely across 

the transition from essentially undeformed rocks on the north 

to substantially deformed rocks to the south, it provides the 

opportunity to document the nature of the early deformational 

phases of the orogeny.

The iron range is part of the Michigamme subterrane of 

the Penokean fold-and-thrust belt terrane, and is the northernmost exposed subterrane of the continental margin assemblage. Rocks of the Michigamme subterrane were deformed 

largely by thin-skinned tectonism, which did not greatly affect 

the Neoarchean basement. Where the Neoarchean basement 

was involved in Penokean deformation, it was by movement of 

fault blocks, rather than by penetrative deformation. The range 

is located within a very foreshortened part of the fold-andthrust belt. The range lies only 50 to 60 kilometers (km) north 

of the Niagara fault, which is the suture between the continental margin and the accreted arcs of the Wisconsin magmatic 

terranes to the south. South of the Gogebic iron range, Paleoproterozoic strata become much more intensely deformed and 

metamorphosed in the subterranes located successively closer 

to the suture zone. 

The Gogebic iron range presents a unique opportunity 

to examine lithologic and structural changes and interpret 

processes and events across a Paleoproterozoic boundary. 

These processes range from stable cratonic sedimentation in 
Introduction 3

E

F

G

A D

C

B

L A K E S U P E R I O R

L

A

K

E

M I C

H I G A

N

L A K E

H

U

R

O

N

Georgian Bay

W I S C O N S I N

M I N N E S O T A

I O W A

M I C H I G

A

N

O N T A R I O

QUEBEC

MANITOBA

O N T A R I O

N i a g a r a F a u l t

G r e n v i l l e

F r o n

CANADA

t

UNITED STATES

0 100 200 300 400 KILOMETERS

95° 90° 85° 80°

50°

45°

EXPLANATION

Sedimentary rocks undivided (Paleozoic)—Limestone, dolomite, sandstone, and shale

Grenville Province (Mesoproterozoic)—Gneiss and plutonic rocks undivided

Midcontinent Rift (Mesoproterozoic)—Flood basalt, rhyolite, sandstone, and gabbroic rocks undivided

Anorogenic pluton (Mesoproterozoic)—Granite and anorthosite undivided

Marquette Range Supergroup and Animikie Group (Paleoproterozoic)—Undivided metasedimentary and metavolcanic rocks.

 Iron ranges shown in black:

 A—Gogebic

 B—Marquette

 C—Menominee

 D—Iron River-Crystal Falls

 E—Cuyuna

 F—Mesabi

 G—Gunflint

Wisconsin Magmatic Terranes (Paleoproterozoic)—Metavolcanic, metasedimentary rocks, and plutons undivided

Huronian Supergroup (Paleoproterozoic)—Sedimentary rocks, gabbroic intrusions undivided

Superior Province (Archean)—Granitic and metavolcanic rocks undivided

Contact

B

Figure 1. Generalized geologic map of the upper Great Lakes region, U.S. and Canada, showing the location of the Gogebic 

iron range and other Paleoproterozoic iron ranges of the Lake Superior region.
4 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

the central part of the range to more tectonically and volcanically active sedimentologic processes in the eastern and 

western parts of the range. This same boundary later formed a 

tectonic front separating the rather intense Penokean deformation in the eastern and western parts of the range from the 

nearly undeformed rocks in the central part of the range. The 

present orientation of the Gogebic iron range also provides a 

unique view of Penokean deformation because of Mesoproterozoic monoclinal tilting. Crustal-scale listric thrusting in the 

interval 1,100 to 1,000 million years ago (Ma) tilted the iron 

range from 40° to 90° northward (Cannon and others, 1993). 

Subsequent erosion resulted in the current map pattern, which 

is essentially a longitudinal cross section of the range as it 

existed after Penokean deformation. 

General Geology

The Gogebic iron range is an arcuate belt of Paleoproterozoic sedimentary and volcanic rocks of the Marquette 

Range Supergroup extending from the western shore of Lake 

Gogebic in the Upper Peninsula of Michigan, westward for 

about 125 km into northern Wisconsin (fig. 1; pl. 1B). The 

range varies from less than 1 km to as much as 5 km wide and 

is bounded by Neoarchean granitic and metavolcanic rocks on 

the south and Mesoproterozoic volcanic and intrusive rocks on 

the north. In gross form, the range is a north-facing monocline 

dipping from about 40o to nearly 90o. In more detail, faulting 

and folding repeat the stratigraphic section in places, particularly in the eastern and western segments. For discussion 

purposes, the range consists of three segments. The central 

segment is characterized by well-defined and laterally continuous stratigraphic units, particularly within the Ironwood 

Iron-Formation, and by a lack of intense Penokean deformational features. In the eastern segment, the detailed internal 

stratigraphy of the Ironwood is lacking, significant changes in 

stratigraphic thickness occur over short intervals, the Ironwood 

interfingers with volcanic rocks of the Emperor Volcanic Complex, and sills of diabase are abundant. Penokean deformational features are also well developed and include folds and 

low-angle faults that have repeated the stratigraphic section. 

The western part of the range lies west of Tyler Forks River 

(pl. 1A). In this part of the range, the internal stratigraphy of 

the Ironwood is obscure, folding and faulting are intense in 

many areas, and large sills of diabase within the Ironwood and 

Palms Formations are common in the westernmost portions 

of the area. The stratigraphic relations within the Marquette 

Range Supergroup and its overlying and underlying units are 

shown in a schematic stratigraphic section in figure 2.

Neoarchean Rocks

Granitic and volcanic rocks of Neoarchean age bound the 

Paleoproterozoic sequence of the Gogebic iron range on the 

south and constitute the basement on which it was deposited 

unconformably. The Neoarchean rocks were best described by 

Schmidt (1976) and Greathead (1975) for part of the central 

Gogebic iron range and by Sims and others (1977) for the 

eastern part of the area. They were also mapped in varying levels of detail by Trent (1973), Prinz and Hubbard (1975), and 

Klasner and others (1998). Except in the easternmost part of 

the area, a scarcity of outcrops has hampered detailed examinations of the Neoarchean rocks. 

A variety of volcanic rock is present and collectively 

forms the Neoarchean Ramsay Formation as defined by 

Schmidt (1976). The Ramsay includes thick sequences of pillowed to massive metabasalt (greenstone), particularly in the 

east. This metabasalt sequence is uniformly south-facing, as 

indicated by the abundant pillows, and dips steeply. When the 

Mesoproterozoic northward tilting is removed, the metabasalt 

sequence is restored to a north-dipping and downward-facing 

section that apparently was part of the overturned limb of a 

recumbent fold that formed during the Neoarchean (Klasner and others, 1998). Other units of the Ramsay Formation 

include fragmental volcanic rocks of mafic to felsic composition, volcanogenic graywackes, and minor argillite. Between 

Hurley and Upson, Wis., the volcanic rocks are at least 6 km 

thick and range in composition from basalt to rhyolite; dacite 

breccia is the most common rock directly beneath the Paleoproterozoic unconformity (Greathead, 1975). In most areas, 

the Ramsay has undergone only greenschist-facies metamorphism; however, locally, especially near the contact with the 

Puritan batholith, the Ramsay has been metamorphosed to 

amphibolite facies. 

Granitic rocks of the Neoarchean Puritan Quartz Monzonite (Schmidt, 1976) form a batholith that intrudes the Ramsay 

Formation. The Puritan Quartz Monzonite (pl. 1B) is mostly 

massive to porphyritic, locally weakly foliated, and contains 

abundant pegmatitic segregations. A Rb-Sr isochron age of 

2,710±140 Ma was determined by Sims and others (1977) 

for massive granitic rocks of the Puritan Quartz Monzonite. 

Strongly gneissic rocks are common, especially near the adjacent volcanic rocks of the Ramsay Formation, and have been 

given local names in several places (Fritts, 1969; Schmidt, 

1976). The locally gneissic units are not distinguished from 

the more massive phases of the batholith on illustrations in this 

chapter because outcrops are insufficient to do so in most of 

the area and because the emphasis of the report is on Paleoproterozoic features.

Paleoproterozoic Rocks

Paleoproterozoic sedimentary and lesser volcanic rocks 

constitute the Marquette Range Supergroup, which includes 

the Ironwood Iron-Formation, the principal iron-bearing strata 

of the Gogebic iron range. The Marquette Range Supergroup 

is composed of three groups that are separated from each other 

by regional unconformities (fig. 2). From oldest to youngest they are: (1) the Chocolay Group, composed of the basal 

Sunday Quartzite and overlying Bad River Dolomite; (2) the 
General Geology 5

Western Gogebic Range Central Gogebic Range Eastern Gogebic Range

Mellen Complex

1102 Ma

Puritan Batholith

2710 Ma

Anvil

Pence

Norrie

Yale

Plymouth

EXPLANATION

Keweenawan Supergroup (Mesoproterozoic)

 Jacobsville Sandstone

 Mellen Intrusive Complex

 Powder Mill Group

Marquette Range Supergroup (Paleoproterozoic)

 Baraga Group

 Tyler Formation (west), Copps Formation (east)

 Metagabbro and metadiabase

 Menominee Group

 Emperor Volcanic Complex

 Ironwood Iron-Formation

 Palms Formation, upper quartzite

 Palms Formation, lower argillite

 Chocolay Group

 Bad River Dolomite

 Sunday Quartzite

Puritan batholith (Neoarchean)

Ramsay Formation (Neoarchean)

Contact

Fault

Unconformity

Figure 2. Schematic stratigraphic section along approximately 150 km of the Gogebic iron range illustrating the stratigraphic 

relationships within the Menominee Group, and the relation of the Menominee Group to other stratigraphic sequences in the region. 

The base of the Menominee Group is used as a horizontal datum. Five members of the Ironwood Iron-Formation are shown in the central 

Gogebic iron range.

Menominee Group, composed of the Palms Formation and the 

overlying Ironwood Iron-Formation, as well as the Emperor 

Volcanic Complex; and (3) the Baraga Group, composed of 

the Tyler Formation in the western and central part of the iron 

range and the laterally equivalent Copps Formation in the 

eastern part of the iron range. Paleoproterozoic intrusive rocks 

(diabase and gabbro dikes and sills) are also common throughout the iron range. Dikes are abundant in the Neoarchean rocks 

and constitute two swarms trending northeast and northwest. 

Dikes are also common in the Chocolay and Menominee 

Groups. These latter dikes were identified mostly in extensive 

underground mine workings and are very seldom seen in natural exposures. Large diabase sills occur mostly in the eastern 

and western extremes of the range and mainly intrude the Ironwood Iron-Formation and Emperor Volcanic Complex. 
6 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

Detailed descriptions of the Marquette Range Supergroup 

strata are given in papers by Irving and Van Hise (1892), Van 

Hise and Leith (1911), Allen and Barrett (1915), Hotchkiss 

(1919), Aldrich (1929), Huber (1959), Alwin (1976), and 

Schmidt (1980). The following brief descriptions are based 

largely on these earlier reports and are supplemented by our 

own observations. This report emphasizes features that help 

explain the contrast between the depositional and tectonic 

styles of the central part of the range and those of the eastern 

and western parts.

Chocolay Group

The Chocolay Group is absent from the central part of 

the Gogebic iron range but is present as more or less continuous strata in both the eastern and western parts. The group 

includes the basal Sunday Quartzite, an orthoquartzite that 

grades upward into dolomite and dolomitic marble of the Bad 

River Dolomite.

Sunday Quartzite

 In the eastern part of the Gogebic iron range, the Sunday 

Quartzite is the basal Paleoproterozoic unit and forms an 

apparently continuous belt about 9 km long (Klasner and 

others, 1998). The eastern part of this outcrop belt is shown 

in figure 3. The lower part of the Sunday is a reddish, prominently crossbedded orthoquartzite that contains conglomeratic 

layers of quartz and granitoid cobbles as large as 8 centimeters (cm) diameter near its base. The rest of the formation is 

a gray, vitreous quartzite with common crossbeds and local 

current ripple marks. Schmidt (1980) reported that the Sunday 

is as much as 46 meters (m) thick, and grades upward into the 

Bad River Dolomite. Perhaps the best exposure of the Sunday 

Quartzite is on the southwest side of the hill in the SW1⁄4 sec. 

18, T. 47 N., R. 44 W. (Klasner and others, 1998), west of the 

area shown in figure 3. Some of the crossbeds are of the herringbone type, where successive crossbeds are oriented 180o 

to each other. These units, along with numerous mud-chip 

conglomerates and mud-cracked units, form a series of stacked 

deposits that are each about 2 m thick. Each sequence starts 

with a mud-cracked unit, is overlain by a herringbone-crossbedded unit, and terminates with a parallel-bedded unit. All of 

these features suggest that the Sunday Quartzite was deposited in a tidal environment (Ojakangas, 1983). The Sunday 

Quartzite is not known in the central and western parts of the 

Gogebic iron range where either the Bad River Dolomite or 

Palms Formation lies directly on Neoarchean basement rocks 

instead.

Bad River Dolomite

The Bad River Dolomite is conformable with and gradational into the Sunday Quartzite in the eastern Gogebic iron 

range. The transition is marked by several meters of interbedded quartzite and dolomite; dolomite beds become thicker 

and more abundant upward. The Bad River contains beds and 

irregular patches of gray to black chert and stromatolitic layers 

and mounds that range from 5 to 50 cm across. In the eastern 

Gogebic iron range, the Bad River weathers to a distinctive 

brown color, suggesting that the dolomite contains a significant amount of iron. Schmidt (1980) reported that the dolomite 

reaches a maximum thickness of about 120 m in the eastern 

Gogebic iron range. The formation thins westward and was 

either eroded or was not deposited from the Wakefield, Mich. 

area westward for a distance of about 75 km.

The Bad River reappears in the western Gogebic iron 

range, west of the Tyler Forks River in Wisconsin (pl. 1A), 

as apparently isolated lenses of dolomite resting on Neoarchean basement; most are too small to show at the current 

map scale. A unit as much as several tens of meters thick of 

variegated chert breccia overlies Neoarchean basement from 

the Tyler Forks River for about 20 km westward (pl. 1A). The 

unit is interpreted to be a residual deposit of chert fragments 

resulting from the weathering and dissolution of the cherty 

dolomite and the accumulation and slight reworking of the 

residuum. Although this unit is included in the Palms Formation (described below) in this report, its presence indicates that 

the Bad River was once extensive in the western part of the 

Gogebic iron range and its absence results from weathering 

and erosion prior to deposition of the Palms Formation. 

Near Mineral Lake (pl. 1A) and near Atkins Lake, Wis., 

at the far western end of the range, the Bad River reaches a 

thickness of about 300 m (Cannon and others, 1996). Near 

Atkins Lake the Bad River consists of gray, dolomitic marble 

containing abundant lenses and layers of chert. Stromatolitic 

beds alternate with thinly laminated or massive dolomite. Most 

of the stromatolites exposed in the abandoned quarry in the 

NW1⁄4 sec. 22, T. 44 N., R. 5 W. (pl. 1A) are low mounds less 

than 25 cm across. However, a large glacial boulder several 

kilometers to the southwest contains silicified stromatolites 

nearly a meter across (fig. 4). 

The basal portion of the Bad River crops out in the SE 

1⁄4 sec. 15, T. 44 N., R. 5 W. (pl. 1A). Here, the Bad River 

is an arkosic quartz-pebble conglomerate that rests directly 

on Neoarchean granitic rocks. The conglomerate grades up 

through about 10 m into arkosic dolomite and eventually into 

the cherty dolomite that constitutes the bulk of the unit. 

Metamorphism in the Atkins Lake area resulted in extensive development of actinolite and tremolite from the reaction 

between the dolomite and chert. The metamorphism occurred 

within the contact aureole of the Mesoproterozoic gabbro of 

the Mineral Lake intrusion.
General Geology 

7

8 9 10 11 12 7 8 9

17 16 15 14 13 18 17 1620 21 22 23 24 19 20 2129 28 27 26 25 30 29 2832 33 34 35 36 31 32 33

2

Creek

Wolf

Mountain

Brotherton

Little

Presque

Isle

River

Wolf

Mountain

Creek

Kinnickinnic Creek

Yondota

Falls

Presque

Isle

River

Ethelwood

Theilers

Ys

Yj

Xi

Xe

Xp

Xp

Xb

Xs

Xmd

Xi

Xe

Xe

Xc

Xe

Ys

Xc

Yj

Ys

Xc Xi

Xp

Wr

Wp

Xe

Xi

Xi

Xi

Wr

Xmd

Xmd

Xi

Xp

Wr

WpSunday Lake fault

Wolf Mountain anticline

Little Presque Isle fault

Presque Isle fault

89° 50’ 89° 48’ 89° 46’ 89° 44’ 89° 42’

48° 28’48° 26’

R. 44 W. R. 43 W.

T. 47 N.

0 1 2 3 KILOMETERS

EXPLANATION

Keweenawan Supergroup (Mesoproterozoic) Jacobsville Sandstone Siemens Creek Volcanics—BasaltMetadiabase (Paleoproterozoic)—Variably metamorphosed diabase and gabbroMarquette Range Supergroup (Paleoproterozoic) Baraga Group Copps Formation—Graywacke

 Menominee Group Emperor Volcanic Complex—Basalt flows and breccia Ironwood Iron-Formation Palms Formation (undivided)—Argillite near base, overlain by quartzite Chocolay Group Bad River Dolomite Sunday Quartzite

Puritan batholith (Neoarchean)Ramsay Formation (Neoarchean)— Agglomerate, pillow basalt, and chlorite schistContactFaultAxial trace of anticline showing direction of plunge of fold axis

Xmd

Yj

Ys

Xc

Xe

Xi

Xp

Xb

Xs

Wp

Wr

Figure 3. Geologic map of part of the eastern Gogebic Range. Generalized from Klasner and others (1998).
8 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

Menominee Group

Strata of the Menominee Group are present throughout 

the western and central parts of the Gogebic iron range and 

are absent only in the eastern half of the eastern segment. The 

group consists of the Palms Formation and the conformably 

overlying Ironwood Iron-Formation. In the eastern part of 

the range, volcanic rocks of the Emperor Volcanic Complex, 

which interfinger with the Ironwood, are included in the 

Menominee Group.

Palms Formation

The Palms Formation unconformably overlies Chocolay 

Group rocks on the eastern and western ends of the Gogebic 

iron range, and overlies Neoarchean rocks in the central part 

of the range. We have shown only two map units within the 

Palms on plate 1A although a regionally consistent internal stratigraphy is present. The principal part of the Palms 

Formation is a coarsening upward succession of argillite near 

the base (the “quartz slate” unit of older literature) grading 

to quartzite near the top (Aldrich, 1929). Ojakangas (1983) 

refined the description, showing that the Palms has a lower 

argillaceous unit, a central mud-silt-sand unit, and an upper 

sand-rich unit; he also stated that the Palms displays a bedding 

pattern and style that suggest deposition in a tidal environment. These distinctive lithologic subdivisions of the Palms 

are readily recognizable throughout the length of the Gogebic 

iron range suggesting that the depositional environment of the 

Palms was exceptionally uniform over the region.

We have defined an additional unit at the base of the 

Palms that occurs throughout a 15-km-long belt in the central 

part of the range, from sec. 1, T. 44 N., R. 2 W., to sec. 16, T. 

44 N., R. 3 W., and in much of the eastern part of the range. 

This basal unit consists of a breccia of variegated angular to 

subrounded clasts of chert as much as 40 cm long in a siliceous matrix. Rounded quartz pebbles and sand grains are 

present locally in the matrix between the chert clasts. This 

rock type is interpreted to be a residuum of chert nodules and 

segmented layers, originally contained in the Bad River Dolomite; the residuum formed as a result of prolonged weathering 

and dissolution of dolomite prior to deposition of the overlying Menominee Group sediments. In the east, the basal breccia 

overlies the Bad River Dolomite. In the central part of the iron 

range, a thin unit of dolomite locally underlies the breccia, 

but most commonly the breccia directly overlies Neoarchean 

basement rocks.

Magnetite and hematite concentrations are present locally 

at the top of the basal chert breccia unit. Shallow exploration 

shafts were excavated into some of these magnetite-rich zones 

during the early days of exploration for iron ore in the region. 

Several such exploration shafts are on the ridge extending 

east from Mount Whittlesey in Wisconsin; several others are 

located 1 km west of where the Bad River transects the unit 

(known informally as Penokee Gap). The presence of these 

magnetite concentrations within sandy chert breccia units on 

an erosion surface suggests that they are paleo-placer deposits 

and that their intermittent occurrence indicates deposition in 

streams. Chemical analyses of several of the deposits failed to 

show any heavy minerals other than magnetite. The magnetite presumably was derived from Neoarchean rocks exposed 

during the erosion interval. The probable paleo-placer deposits 

support the interpretation of Klasner and LaBerge (1996) that 

the chert breccia is a lag accumulation on an erosion surface. 

Although the regional unconformity developed on the Bad 

River Dolomite has been known since the work of Irving and 

Van Hise (1892), a recent interpretation by Sims and others 

(1989) suggests that the erosion was related to slight rotation 

of blocks of Neoarchean basement along the Penokean continental margin.

Ironwood Iron-Formation

The Palms Formation grades upward over an interval of 

several meters into the Ironwood Iron-Formation, marking a 

change from clastic to chemical sedimentation. The Ironwood, 

like iron-formations elsewhere in the Lake Superior region, 

has two distinct lithologies and bedding styles, as described by 

Huber (1959): (1) “wavy-bedded ferruginous chert” consisting 

of irregular beds and lenses of granular or oolitic chert (a few 

millimeters to several centimeters thick) which are separated 

by thin laminae of more evenly bedded, iron-rich material; 

the iron-rich laminae consist largely of iron oxides, silicates 

(mostly minnesotaite and stilpnomelane), and carbonate; and 

(2) “even-bedded ferruginous slate” in which the beds are 

extremely regular and finely laminated. The chert is generally 

dense and flinty. These two lithologies are generally thought 

to reflect different depths of water in which the iron-formation 

accumulated; the wavy-bedded iron-formation was deposited 

in shallow (agitated) water, and the laminated iron-formation 

was deposited in deeper (quieter) water. 

Figure 4. Stromatolites in Bad River Dolomite. Photograph is of 

a large boulder in NE1⁄4, sec. 5. T. 43 N., R. 5 W. Field of view about 

0.5 m across. 
General Geology 9

Based on a preponderance of one or the other of these 

different bedding styles, Hotchkiss (1919) divided the Ironwood within the previously productive central part of the iron 

range into five members. These are, from the base upward, the 

Plymouth, Yale, Norrie, Pence, and Anvil Members (fig. 2). 

The Plymouth Member is 40 to 46 m thick and consists mainly 

of irregularly bedded, granular iron-formation containing thin 

intraformational conglomerates. Stromatolitic units are present 

locally. The overlying Yale Member is 14 to 22 m thick, has 

a 3-m-thick basal unit of pyritic shale (suggesting periodic 

anoxic conditions), and consists mainly of laminated chertcarbonate iron-formation and some granular or intraformational conglomerate units. Overlying the Yale Member is the 

Norrie Member which is 34 to 36 m thick and consists mainly 

of irregularly bedded, granular iron-formation. The 30- to 

37-m-thick Pence Member consists mainly of laminated chertcarbonate-silicate iron-formation containing some irregularly 

bedded, granular units. The uppermost unit, the Anvil Member, is a 16-m-thick, magnetite-rich, laminated chert-carbonate 

iron-formation containing some irregularly bedded, granular 

horizons. Additional details of the stratigraphic subdivisions of 

the Ironwood were provided by Huber (1959).

The central part of the range has few natural outcrops so 

our understanding of it depends largely on past studies such 

as by Hotchkiss (1919), Aldrich (1929), and Huber (1959), 

that were conducted when extensive underground mining 

and exploration drilling was taking place, and from the more 

recent summary by Schmidt (1980). Few new observations of 

the Ironwood were made in the central part of the iron range 

during this study. The older data show that the five members 

of the Ironwood are persistent in character, although of variable thickness, from near Wakefield, Mich., westward for 

about 70 km to the Tyler Forks River west of Upson, Wis. 

Except for the presence of some ferruginous argillite horizons 

and pyritic shale units, the Ironwood contains little detrital 

or volcanic material throughout this belt. It should be noted 

that many authors of early reports on the Gogebic iron range 

used the term “slate” for thinly laminated or fissile, finegrained rocks, even where cleavage is completely lacking. The 

common usage of slate as a lithologic term for much of the 

Gogebic iron range is misleading because most of these rocks 

do not possess a penetrative deformational fabric, but rather 

are simply finely laminated sedimentary rocks. 

East of Wakefield, Mich., the stratigraphy of the Ironwood changes abruptly. A unit of ferruginous argillite from 

30 to 50 m thick that appears within the lower part of the 

Ironwood immediately east of the Sunday Lake fault indicates 

the deposition of clastic material that is absent in the central 

part of the range (Huber, 1959). More strikingly, the Emperor 

Volcanic Complex (Trent, 1976) is interfingered with the 

iron-formation in and near the Presque Isle trough. In addition 

to the volcanic rocks, the Ironwood Iron-Formation contains 

significant detrital material in the eastern Gogebic iron range. 

Allen and Barrett (1915) reported that an argillite unit up to 

147 m thick is within the Ironwood in secs. 16, 17 and 18, T. 

47 N., R. 44 W. An argillite unit of unknown thickness interbedded within the Ironwood is also exposed beneath a mafic 

sill on the west side of the hill in the NW1⁄4 sec. 26, T. 47 N., 

R. 44 W.

Both the abrupt changes in the stratigraphy of the Ironwood east of Wakefield, and the rift-related volcanic activity 

in a fault-bounded basin that included a significant influx of 

detrital material document a tectonically unstable environment 

during deposition of the Ironwood in that area. Inasmuch as 

the underlying Palms Formation retains its internal lithologic 

subdivisions throughout the area, the rifting, subsidence, and 

associated volcanism apparently did not begin until the time of 

deposition of the Ironwood Iron-Formation.

Relation between the Ironwood Iron-Formation and the 

Emperor Volcanic Complex

Volcanic rocks and mafic sills of the Emperor Volcanic 

Complex (Trent, 1976) are a significant component of the 

stratigraphic section in the eastern Gogebic iron range. Furthermore, the various subdivisions of the Ironwood, developed 

so prominently in the central part of the range, are not as 

clearly recognized to the east. Even at mines such as the Sunday Lake and Vicar within a few kilometers east of the Sunday 

Lake fault, the correlation of the internal stratigraphy of the 

Ironwood was controversial among mining geologists (Huber, 

1959). Volcanic rocks of the Emperor Volcanic Complex intertongue with the Ironwood and provide proof that the Ironwood 

was deposited at the same time that volcanism was occurring 

in this region. The Little Presque Isle fault (fig. 3) marks a 

prominent change in the lithology of the Ironwood and in the 

thickness of the Emperor. West of the fault, the Ironwood is 

mostly irregularly bedded, granular iron-formation typical 

of shallow-water deposits. East of the fault, the Ironwood is 

mainly the laminated variety typical of deeper water deposits. 

The Little Presque Isle fault also marks a prominent change in 

the distribution and thickness of the Emperor Volcanic Complex. West of the fault, the Emperor consists of a mafic sill and 

several thin basalt flows interbedded with the Ironwood. East 

of the fault, the Emperor thickens abruptly to as much as 2,000 

m and is composed dominantly of hyaloclastites, pillow breccias, and related debris flows (fig. 5) as well as lesser pillowed 

basalts, crackle breccias, and massive basalt. 

Exposures on the hill in the NE1⁄4 sec. 24, T. 47 N., R. 44 

W. (fig. 4), are representative examples of the rock types in the 

Emperor. A zone containing 2- to 3-m-long pillows is overlain by a debris flow that has a matrix of 0.5- to 2.0-cm-long 

hyaloclastite fragments in which both pillow fragments and 

complete pillows (commonly containing quench cracks) are 

suspended. Scattered pods and patches of quartz and jasper are 

also present in the debris flows. Many of the pillow fragments 

and most of the hyaloclastite fragments are composed primarily of zoisite. Larger pillow fragments may have a zoisite rind 

and an epidote core, which suggests leaching of iron from 

the fragments. The deposits are interpreted to be debris that 

flowed off a subaqueous lava dome (or domes). The abrupt 

change in stratigraphy across the Little Presque Isle fault sug-
10 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

gests that it was a growth fault along the northwest side of a 

graben that formed during Ironwood deposition. Farther east, 

the Emperor thins toward the Presque Isle fault, which suggests that the fault bounds the southeastern side of the graben 

(fig. 3) (Klasner and others, 1998) that is herein named the 

Presque Isle trough.

Chemically, the Emperor consists of rift-related tholeiite 

(Sims and other, 1989) and mafic rocks, all of which form the 

vast majority of the volcanic sequence; however, subordinate 

amounts of volcanic rocks composed mainly of fine sericite 

with well-preserved shard structures and quartz phenocrysts 

suggest that felsic volcanic rocks are at least a minor component. Dan (1978) and Prinz (1981) report finding rocks as 

felsic in composition as dacite in the Emperor. More details 

of the chemistry of the Emperor Volcanic Complex are given 

below in the section on “Paleoproterozoic Intrusive Rocks.”

Stratigraphic Relations in the Menominee Group and 

Emperor Volcanic Complex

Both the abrupt changes in the stratigraphy of the Ironwood Iron-Formation east of Wakefield, Mich., and the riftrelated volcanic activity in a fault-bounded basin that includes 

a significant influx of detrital material document a tectonically unstable environment of deposition of the Ironwood in 

that area. Because the underlying Palms Formation retains its 

internal lithologic subdivisions throughout the area, the rifting, 

subsidence, and associated volcanism apparently did not begin 

until the time of Ironwood deposition.

Similar stratigraphic relations within the Ironwood IronFormation are present on the western end of the Gogebic iron 

range. Details of the geology of the western part of the range 

are shown on plate 1A. The five-fold internal stratigraphy of 

Figure 5. Examples of typical rocks types of the Emperor Volcanic Complex. A, Debris flow containing pillow breccia fragments in a 

finer-grained hyaloclastite matrix from outcrop near Wolf Mountain. B, Polished surface of debris flow breccia. Paper clip is 3 cm long. 

C, Photomicrograph of hyaloclastite breccia.D. Photomicrograph showing perlitic cracks within a clast contained in debris flow.

1 centimeter

A B

C D
General Geology 11

the Ironwood in the central part of the range cannot be traced 

west of the village of Tyler Forks in sec. 32, T. 45 N., R. 1 

W. The difficulty in tracing members of the Ironwood may 

result (in part) from an increased metamorphic grade toward 

the west within the contact aureole of the Mellen Intrusive 

Complex (Mesoproterozoic), and (in part) from an increasing 

complexity of internal folding and faulting in the Ironwood. In 

this study the various members could not be recognized in the 

western segment, even in the least metamorphosed and least 

deformed sections. The five-fold subdivision of the central part 

of the range probably does, indeed, lose its definition toward 

the west, perhaps as a result of somewhat more variable 

depositional environments in the western part of the range. 

Although the detailed studies conducted by U.S. Steel Corporation in the 1950s distinguished internal units in the Ironwood 

as far west as Penokee Gap on the Bad River (summarized by 

Marsden, 1978), the definition of the units was based, in part, 

on metallurgical characteristics. They did not correlate entirely 

with the classic five-member subdivision of the central part of 

the range.

Outcrops in the NE1⁄4 sec. 20, T. 44 N., R. 5 W., in the 

Atkins Lake area (pl. 1A) contain both the laminated and 

granular varieties of the Ironwood Iron-Formation, which 

overlie at least 15 m of thick-bedded quartzite that is typical of 

the upper Palms Formation. The change in lithology is further 

displayed by laminated argillite as much as 15 m thick that 

forms interbeds in the lower part of the Ironwood. The argillite is succeeded by a unit of amygdaloidal basalt about 6 m 

thick containing numerous tabular to round, randomly oriented 

blocks of argillite and iron-formation as much as several 

meters long. Some of the fragments are relatively rounded and 

are internally intensely deformed; others are sharply angular 

and have little or no internal deformation. The basalt matrix 

adjacent to some of the larger fragments is discolored and 

contains more vesicles than is typical, which suggests that the 

basalt may have chilled against the fragments. A meter-thick 

zone of laminated, magnetic iron-formation (which may be 

a large slab included in the basalt) separates the basalt from 

an overlying, 10- to 20-m- thick metadiabase sill. The upper 

margin of the sill is chilled against a breccia unit containing 

chert-rich argillite clasts in an amygdaloidal basalt matrix 

similar to the unit beneath the sill. Local alteration rinds that 

are suggestive of pillows or pillow breccia are present in the 

basalt. Outcrops in sec. 20, T. 44 N., R. 5 W., have zones of 

hyaloclastite-bearing basalt.

At least three metadiabase sills are present in the Ironwood Iron-Formation in the Atkins Lake area. Their intrusion into the semi-consolidated argillite and iron-formation 

produced peperite-like features caused by the interaction of the 

basaltic magma with the wet sediments (Klasner and LaBerge, 

1996). The peperite textures suggest that the sill emplacement 

was nearly contemporaneous with the iron-formation deposition. Several drill cores from near Trapper Lake (sec. 26, T. 44 

N., R. 6 W.) (pl. 1A) about 5 km west along strike from Atkins 

Lake, intersected laminated and irregularly bedded ironformation and interbedded turbidites, a 70-m-thick graphitic 

argillite, and several metadiabase sills that had highly brecciated border zones similar to those exposed near Atkins Lake.

Collectively, these data suggest that the depositional 

environment of the Ironwood changed westward from a stable 

platform environment into an unstable basin. Igneous activity, turbidites, and at least periodic stagnant anoxic conditions 

accompanied iron-formation deposition. A comparison of 

the lithologic changes with those of the eastern Gogebic iron 

range suggests that the igneous activity and turbidites in the 

west may also have been a response to the formation of faultbounded depositional basins. The presence of such structures, 

however, is not obvious within the present exposures. The 

Palms Formation retains its lithologic units and thickness both 

in the eastern Gogebic iron range and throughout the area, 

indicating that the unstable conditions and volcanism developed contemporaneously with the deposition of the Ironwood 

Iron-Formation. 

Although volcanism that is contemporaneous with the 

deposition of iron-formation is well represented on the eastern and western ends of the range, volcanism has not been 

reported in the central part of the range where extensive mine 

openings and drilling have exposed essentially continuous 

sections of the Ironwood Iron-Formation. However, Paleoproterozoic metadiabase dikes and sills are common in the central 

part of the range where numerous dikes have cut the ironformation at nearly right angles to bedding. Less commonly, 

sills were emplaced along the bedding planes. These rocks are 

discussed in more detail below in the section on Paleoproterozoic intrusive rocks. The dikes were, in fact, important features 

in controlling the location and distribution of iron- ore bodies 

in the productive central part of the range. The age of these 

dikes is not known other than that they are somewhat younger 

than the Ironwood Iron-Formation and possibly older than the 

Tyler Formation. These dikes may have been feeders for volcanic rocks higher in the section, but if so, such volcanic rocks 

have not been recognized. The dikes do not seem to be present 

in the overlying Tyler Formation. Therefore, the available 

evidence suggests that there was igneous activity over much of 

the length of the Gogebic iron range more or less at the same 

time as iron-formation deposition.

Baraga Group

The Baraga Group, which contains the youngest Paleoproterozoic rocks in the Gogebic iron range, is a thick 

sequence of dominantly turbiditic rocks. These rocks overlie 

the Menominee Group and display a contact that varies from 

a definite unconformity in the east to a more problematic, but 

probably a low-angle unconformity or disconformity in the 

central and western parts of the range. This greywacke and 

slate sequence was assigned different names in various areas 

of Wisconsin and Michigan. In the central and western parts 

of the Gogebic iron range, it was named the Tyler Formation 

by Van Hise (1901). Rocks that are very similar lithologically 

and stratigraphically in the eastern Gogebic range were named 

Copps Formation by Allen and Barrett (1915). Although there 
12 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

is little doubt that these two units are correlative, they are 

physically separated by only a few kilometers across an area 

where the Baraga Group has been removed by erosion, and 

thus they continue to maintain separate names. The sequence 

is thickest in the eastern and western parts of the range, but 

was completely removed by erosion prior to the deposition of 

the Keweenawan Supergroup (Mesoproterozoic) near Wakefield, Mich. (pl. 1B). The Tyler and Copps Formations are 

generally correlated with the Michigamme Formation, which 

is widely distributed in other parts of northern Michigan.

Copps Formation

The Copps Formation crops out from near Wakefield, 

Mich., eastward to Lake Gogebic (fig. 3, pl. 1B). The Copps 

overlies the Ironwood Iron-Formation and Emperor Volcanic 

Complex in the western part of its outcrop area, and overlies 

Neoarchean granitoid rocks to the east toward Lake Gogebic. 

The Copps clearly overlies, with an unconformable contact, 

the older rocks in the eastern part of the Gogebic iron range, as 

indicated by the progressive eastward truncation of pre-Copps 

units (Klasner and others, 1998).

The basal part of the Copps is a brown-weathering 

quartzite about 5 m thick that contains several conglomeratic 

zones. Conglomerates near the Presque Isle River in sec. 20, 

T. 47 N, R. 43 W. (fig. 4) contain abundant chert and jasper 

clasts, along with some clasts of volcanic rocks, all of which 

are similar to the rocks of the underlying Ironwood IronFormation and Emperor Volcanic Complex. To the east, in 

secs. 23 and 24, T. 47 N., R. 43 W., where the Copps directly 

overlies Neoarchean granitoid rocks, the basal conglomerate contains abundant granitoid boulders as large as 40 cm in 

diameter.

The basal quartzite unit is overlain by a thick greywacke 

and slate sequence containing abundant graded beds. Sandsized grains in some graded units consist of quartz, plagioclase, microcline, recrystallized chert, and sericitized volcanic 

rocks, all of which indicate a source area containing granitoid, 

sedimentary, and volcanic rocks. The coarser parts of the 

graded beds also commonly contain carbonate concretions. 

Individual turbidite units, which are generally no more than 30 

cm thick, locally contain conglomeratic zones and are consistently north-facing. 

Tyler Formation

The Tyler Formation is present from Wakefield, Mich., 

westward (pl. 1B), and its preserved thickness increases to 

more than 3,000 m in Wisconsin. The Tyler is primarily a 

turbidite succession of intercalated graywacke and black slate, 

argillite, and chert containing abundant Bouma sequences 

(Alwin, 1976). Grains in the graywacke suggest a source area 

with a significant granitic component. Paleocurrent indicators 

suggest transport to the northwest (Alwin, 1976). The lower 

part of the Tyler, including as much as 100 m of strata known 

mainly from exploration drill holes and mine workings, consists of sideritic black slate and ferruginous conglomerates.

The nature and exact stratigraphic position of the Tyler’s 

contact with the underlying Ironwood Iron-Formation has 

been in doubt for many years and various authors have argued 

for either a conformable or unconformable contact. Because 

the only exposures of the basal part of the Tyler were in mine 

workings (now long-abandoned) and in exploration drill 

holes, there has been no recent evidence to resolve this issue. 

A ferruginous (sideritic) shale unit which is as much as 90 m 

thick and contains as much as 25 percent iron, is present in the 

lower part of the Tyler Formation (Hotchkiss, 1919). Atwater 

(1938) stated that the ferruginous shale unit is present throughout most or all of the Gogebic iron range, and that a conglomeratic quartzite (the Pabst Member described later) forms the 

basal part of the Tyler Formation in the eastern part of the 

range. Siderite-chert iron-formation has also been widely 

reported in the basal 100 m of the Tyler (Hotchkiss, 1919). 

Pabst Member

Van Hise (1901) named the Tyler Formation and considered it to conformably overlie the Ironwood. However, 

Hotchkiss (1919, p. 18) suggested that the Pabst Member, a 

conglomerate with pebbles of iron oxide, quartz and jasper, 

was the basal unit of the Tyler Formation, and that the Pabst 

was deposited unconformably over the Ironwood. Hotchkiss’ 

data came mainly from exposures in mine workings and drill 

cores in the productive central part of the Gogebic iron range. 

Aldrich (1929, p. 165–166) concurred with Irving and Van 

Hise (1892) that the Tyler conformably overlies the Ironwood. 

Atwater (1938) examined the relationship between the Ironwood and the Pabst Member as it is exposed in a number of 

mines between Iron Belt, Wis., and Wakefield, Mich., and concluded that an erosional surface existed between the Ironwood 

and Tyler Formations. His diagram (Atwater, 1938, p. 163), 

used as a basis for figure 6, shows that a substantial part of the 

upper Ironwood was eroded prior to deposition of the Pabst. 

Descriptions of the Pabst Member from Hotchkiss (1919) and 

Atwater (1938) suggest that it is similar to the conglomeratic quartzite that forms the base of the Copps Formation in 

the eastern part of the Gogebic iron range; however, Aldrich 

(1929) reported that a shaley unit forms the base of the Tyler, 

and that it underlies the ferruginous shale unit on the western 

end of the range, where an unconformity between the Ironwood and Tyler is uncertain. Finally, Schmidt and Hubbard 

(1972) and Schmidt (1980), based on a review of this older 

data, concluded that the contact between the Ironwood and 

Tyler Formations is most likely gradational and that sedimentation was continuous in the Bessemer, Mich. area. They state 

that the Ironwood grades up into a shale and that, in the Bessemer area, the Pabst Member occurs above the shale. 

Based solely on the evidence of past reports and on 

regional stratigraphic relationships, we believe that the base of 

the Tyler Formation should be placed at the base of the Pabst 

Member and that a low-angle to disconformable contact exists 

between the Pabst and Ironwood. The findings of Atwater 

(1938) that the Pabst truncates the internal units of the Iron-
General Geology 13

wood at a low angle seems to be compelling evidence of the 

unconformable relationship. The Ironwood may indeed have 

graded up into a shale as suggested by Schmidt and Hubbard (1972) and Schmidt (1980), but this shale is preserved 

only locally, as at Bessemer, where pre-Pabst erosion did not 

cut down into the Ironwood Iron-Formation. The interpretation that the Pabst is the basal conglomerate of the Baraga 

Group and that the group is unconformable on the underlying 

Ironwood Iron-Formation is fully consistent with the relation 

of the Copps to older units in the eastern part of the iron range 

and with relationships elsewhere in northern Michigan where 

an unconformity between the Baraga Group and older rocks 

is well established. We propose, therefore, to establish the 

Pabst Conglomerate as the basal member of the Tyler Formation. Further, we consider the relatively local units of black 

sideritic shale that underlie the Pabst Member as upper beds of 

the Ironwood Iron-Formation and mark a transition from the 

relatively shallow-water, oxygenated depositional environment 

of the Anvil Member to euxinic conditions marking the close 

of the iron-formation deposition in this area.

Age and Correlation of the Marquette Range 

Supergroup

Stratigraphic Correlations within the Marquette Range 

Supergroup

The various lithologic units that compose the Marquette 

Range Supergroup in the Gogebic iron range are part of a 

more geographically extensive rock sequence that encompasses the other iron ranges of northern Michigan and Wisconsin (Marquette, Menominee, Iron River-Crystal Falls iron 

ranges) and the intervening areas of non-iron-bearing strata. 

Modern stratigraphic correlation for these rocks dates from 

the work of James (1958) who introduced and defined the 

group names still in use (Chocolay, Menominee, and Baraga 

Groups for the Gogebic iron range). James (1958) also 

introduced the term “Animikie Series” for the entire Paleoproterozoic sequence, implying an approximate correlation with 

the Animikie Group of Minnesota. Cannon and Gair (1970) 

introduced the name “Marquette Range Supergroup” in place 

of “Animikie Series” in compliance with the North American 

Stratigraphic Code (several groups constitute a supergroup, not 

a series, and “Animikie” was already established as a group 

name).

With regard to the Gogebic iron range, until recently, the 

only modification of the stratigraphic terminology defined 

by James (1958) was the formal designation of Emperor 

Volcanic Complex by Trent (1976). The Emperor, as therein 

defined, is a volcanic unit of formation rank and is part of 

the Menominee Group. Trent recognized its essential contemporaneous and intertonguing relation with the Ironwood 

Iron-Formation. Recently, an alternative correlation of some 

units in the Gogebic iron range was proposed by Morey (1996) 

and Ojakangas and others (2001). These authors proposed to 

reclassify the Palms Formation, Ironwood Iron-Formation, and 

Emperor Volcanic Complex as part of the Baraga Group rather 

than accepting their long-accepted placement in the Menominee Group. The correlation that they proposed accepted the 

Chocolay Group, as previously defined, to include the Sunday 

Quartzite and Bad River Dolomite. They proposed that the 

Menominee Group is completely absent in the Gogebic iron 

range. They thus correlated the Palms Formation with units 

such as the Goodrich Quartzite in the Marquette iron range 

and the Ironwood Iron-Formation with the Bijiki Iron-Formation Member of the Michigamme Formation in the western 

Marquette iron range. 

We do not accept this new interpretation for several 

reasons. Although Morey (1996) and Ojakangas and others (2001) acknowledge that an unconformity separates the 

Copps Formation from the Emperor Volcanic Complex in 

the eastern Gogebic iron range, they also accept, without 

question, a conformable contact between the Tyler and the 

Ironwood farther west as proposed by some previous authors. 

As discussed previously, the nature of that contact is still not 

firmly established because exposures are no longer accessible; 

however, we believe that a preponderance of evidence favors 

an unconformable contact between the basal part of the Tyler 

(its Pabst Member) and the underlying Ironwood Iron-Formation, especially based on the evidence that the Pabst transects 

stratigraphic units in the Ironwood. Morey (1996) and Ojakan- WISCONSIN MICHIGAN

Upper unnamed shale member

Anvil Member

Pence Member

Pabst Member of Tyler Formation

0 1 2 3 KILOMETERS

75

50

25

0

25

METERS

75

50

25

0

25

METERS

Figure 6. Schematic longitudinal section of part of the central Gogebic iron range illustrating the crosscutting relations between 

the Pabst Member of the Tyler Formation and underlying members of the Ironwood Iron-Formation. Simplified from Atwater (1937).
14 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

Figure 7. Geochemical discrimination diagrams showing the subalkaline tholeittic character of the mafic rocks from the Emperor 

Volcanic Complex (solid squares) and diabase near Atkins Lake. See table 1 for details of each sample. Boundaries in A from 

Winchester and Floyd (1977). Tholietic-calc-alkaline boundary in B after Irvine and Baragar (1971).

A B

EXPLANATION

Emperor Volcanic Complex

Diabase near Atkins Lake, 

 Wisconsin Tholeiitic

Calc-alkaline

FeO(t)

Na2O + K2O MgO

Rhyolite

Rhyodacite/dacite

Andesite

Andesite/basalt

Sub-alkaline basalt

Alkaline basalt

Trachyandesite

0.001

0.01

0.1

1

0.01 0.1 1 10

Nb/Y

Zr/TiO2

gas and others (2001) cite the presumed conformable contact 

as the principal evidence that the Palms, Ironwood, and Tyler 

units are a continuously deposited sequence and are therefore 

correlative with the Goodrich, Michigamme, and Bijiki succession of the Marquette iron range. 

We believe that the clear evidence of an angular unconformity at the base of the Copps Formation, and at least the 

possibility of an unconformity at the base of the Tyler Formation both argue against the new correlation by Morey (1996) 

and Ojakangas and other (2001). In the western Marquette 

iron range, the base of the Baraga Group is represented by 

the Goodrich Quartzite, which includes a substantial thickness of conglomerate containing iron-formation detritus. This 

conglomerate is therefore equivalent to the basal conglomerate of the Copps Formation described above and to the Pabst 

Member of the Tyler Formation as well. 

The Goodrich Quartzite grades upward into graphitic 

and pyritic black slate of the lower part of the Michigamme 

Formation. The overlying Bijiki Iron-Formation Member of 

the Michigamme Formation is a thin (100 m or less) carbonate 

and silicate iron-formation of local importance in the western 

part of the Marquette iron range and can be traced for about 

15 km along strike (Cannon and Klasner, 1977, 1978). The 

Bijiki is evenly bedded and apparently was deposited in deep 

and (or) quiet water. The unit is overlain by additional pyritic 

and graphitic black slate of the Michigamme. The Bijiki is 

evidently a product of local iron-rich deposition in a stagnant 

basin and thus is quite distinct from the Ironwood Iron-Formation in which abundant shallow-water and well-oxygenated 

facies are present. The Bijiki is, therefore, much more akin to 

the sideritic shale units that overlie the Pabst Member in the 

lower part of the Tyler Formation. Drilling records indicate 

the sideritic shale of the lower Tyler is much more extensive 

than the Bijiki, but the sideritic shale was never distinguished 

as a mappable unit because of a lack of outcrops. Also, there 

is clear evidence that strong tectonic and volcanic activity 

accompanied the deposition of the Ironwood in contrast to the 

quiet conditions indicated for the Bijiki. Although the true correlation may inevitably rely on future precise geochronological data, the evidence in hand does not warrant the substantial 

recorrelation proposed by Morey (1996) and Ojakangas and 

others (2001). Rather, the available evidence substantially supports the long-accepted correlations that we presented above.

Absolute Age of the Marquette Range Supergroup

Precise, radiometrically determined ages for deposition 

of the Marquette Range Supergroup were not available until 

recent work by Schneider and others (2002). Until then, the 

supergroup was restricted to the approximate interval between 

2.6 Ga (the age of youngest basement rocks) and 1.85 Ga (the 

approximate age of deformation and metamorphism). Schneider and others (2002) reported an age of 1,874±9 Ma based 

on sensitive high-resolution ion microprobe (SHRIMP) U-Pb 

dating of zircons from the Hemlock Formation, a unit about 

50 km east of the eastern part of the Gogebic iron range. The 

Hemlock is a lateral equivalent of the Negaunee Iron-Formation and is also probably equivalent to the Emperor Volcanic 

Complex. A SHRIMP U-Pb zircon age of 1,878±2 Ma for a 

thin volcanic unit in the Gunflint Iron-Formation in Ontario 

was determined by Fralick and others (1998, 2002). The 

Gunflint is widely accepted as a correlative of the Ironwood 

Iron-Formation of the Gogebic iron range. These two ages, 

although determined for rocks somewhat removed from the 

Gogebic range, provide a good approximation for the age of 

deposition of the Menominee Group. Compressional deformation of the Marquette Range Supergroup probably began 

by at least 1,850 Ma and a post-tectonic pluton was dated at 
General Geology 15

1,835±9 Ma (Schneider and others, 2002). The age range of 

Baraga Group deposition is therefore constrained to the interval from about 1,875 to 1,850 Ma.

The age of the Chocolay Group remains more problematic. Chocolay rocks are preserved only sporadically within 

the southern Lake Superior region and lie beneath a regional 

unconformity marking the base of the Menominee Group. 

The Chocolay may, therefore, be substantially older than the 

Menominee Group. Numerous authors have suggested that 

the Chocolay is correlative with the upper parts of the Huronian Supergroup in Ontario, with which it shares substantial 

lithologic similarity, particularly the common occurrence of 

glaciogenic units (not present in the Gogebic iron range) and 

thick orthoquartzite sequences (Puffett, 1969; Young, 1970, 

1973, 1983; Ojakangas, 1984, 1985, 1988; Roscoe and Card, 

1993). If that correlation is correct, the Chocolay Group must 

be older than 2,200 Ma, which is the age of the Nipissing Diabase that cuts the Huronian Supergroup (Corfu and Andrews, 

1986).

Paleoproterozoic Intrusive Rocks

Paleoproterozoic dikes and sills of diabase and gabbro cut 

both the Marquette Range Supergroup and the Archean basement rocks. Sills intrusive into the Ironwood Iron-Formation 

are found on both the eastern and western ends of the Gogebic 

iron range. In the east, several large gabbroic sills (as much 

as 300 m thick) are concordant with the Ironwood and have 

been folded with it. Likewise, in the west, at least three sills 

have intruded the Ironwood. In the intervening central part of 

the range, including the productive mining district, sills are 

sparse but diabase dikes are very common and were probably intruded during the same magmatic event as the sills. 

Unpublished mine maps and sections show approximately 

100 mapped dikes cutting the Ironwood Iron-Formation 

between Wakefield, Mich., and Iron Belt, Wis. The dikes cut 

the Ironwood bedding at approximately right angles and form 

both northeast- and northwest-trending sets. Similar dikes 

intruded the Neoarchean rocks south of the Gogebic iron 

range and were mapped in detail where outcrops are adequate 

(Trent, 1973; Prinz and Hubbard, 1975; Klasner and others, 

1998). These dikes occur in two sets that intersect at a nearly 

right angle; one set trends east-northeast and the other northnorthwest, and all generally dip moderately to the south. The 

attitude of the dikes has been affected by the northward tilting 

of the region during events related to the Midcontinent Rift 

at about 1.1 Ga (Cannon and others, 1993); before the rifting 

event, the dikes were essentially vertical.

Paleoproterozoic dikes are very abundant in rocks of the 

Menominee Group and Neoarchean basement, but they are 

apparently absent in the Tyler and Copps Formations. Areas 

underlain by the Tyler and Copps have only scattered outcrops 

and there is a possibility that dikes are present (but concealed). 

However, some of those areas have outcrop density comparable to areas of Neoarchean rocks where dikes can be readily 

mapped. Diabase dikes are typically more resistant to weathering than the graywacke and shale units of the Tyler and 

Copps and would likely form ridges if present. For example, 

topographic ridges are commonly formed by Mesoproterozoic 

dikes where they have cut the Tyler Formation such as sec. 20, 

T. 45 N., R. 1 W. (pl. 1A). Thus, the mafic intrusive activity 

most likely occurred shortly after deposition of the Menominee Group or possibly concurrent with it, and the igneous 

rocks, like their sedimentary host rocks, probably are truncated by the unconformity between the top of the Menominee 

Group and the base of the Baraga Group. Because the sills 

are approximately coextensive with the Emperor Volcanic 

Complex in the east and are also located near volcanic units 

in the west, the intrusive and volcanic rocks are probably 

comagmatic.

Four samples of mafic volcanic rocks from the Emperor 

Volcanic Complex in the eastern Gogebic iron range were analyzed for major and trace elements. In addition, two samples 

from a mafic sill that intruded the Ironwood Iron-Formation 

near Atkins Lake, Wis., at the western end of the Gogebic iron 

range were analyzed. These geochemical data are presented in 

table 1. 

All samples are subalkaline tholeiitic basalts (fig. 7). 

Sample EMP 7 is less evolved than the other samples as it 

has a higher Mg number (58.7), lower FeOt and TiO2 concentrations, and higher Cr and Ni contents (table 1). The other 

samples have very similar evolved compositions: low Mg 

numbers (48.6–40.2), high FeOt (11.76–13.72 weight percent), 

and TiO2 (1.31–1.87 weight percent). As a group, the samples 

define a trend of increasing FeOt, TiO2, and incompatible 

trace elements (for example, Zr, Nb, Y, La) and decreasing Mg 

numbers typical of tholeiitic basalt suites. Decreasing Cr and 

Ni and decreasing Mg numbers are consistent with fractionation of olivine and clinopyroxene.

Chondrite-normalized rare-earth-element (REE) plots 

indicate that all samples have been enriched in light REE, 

having abundances from 30 to 75 times chondritic abundances 

and heavy REE from 9 to 15 times chondritic abundances (fig. 

8A). None of the samples has a europium anomaly, which suggests limited plagioclase fractionation. Chondrite-normalized 

La/Yb ratios range from 3.3 to 6.6, generally increasing with 

decreasing Mg number (table 1). Compared to the composition 

of primitive mantle, these basalts have variably negative Nb, 

Ta, Ti, and a positive Th anomaly (fig. 8B). The compositional 

similarity between the samples of Emperor basalt and samples 

of diabase from the sill near Atkins Lake supports the correlation of these mafic igneous rocks.

The variation in the La/Yb ratio shown by the basalt 

samples is greater than would result from simple crystal fractionation (Ueng and others, 1988). The variation suggests the 

possible involvement of more complex magmatic processes 

such as crustal assimilation and fractional crystallization. The 

role of that process in the evolution of the Emperor basalts is 

supported by (1) the general correlation of the increasing La/

Yb ratio with the decreasing Mg number, and (2) the variably 

negative Nb, Ta, Ti, and positive Th anomalies on the primitive-mantle-normalized plot (fig. 8B) (Ueng and others, 1988). 
16 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

Table 1. Whole-rock chemical analyses of Paleoproterozoic mafic rocks from the Emperor Volcanic Complex and diabase near Atkins 

Lake.—Continued

[All measurements in weight percent unless otherwise noted. Major elements recalculated to 100-percent volatile free. Samples 1–4 from Emperor Volcanic 

Complex; samples 5 and 6 from diabase near Atkins Lake. Fe0t is total iron as Fe0. LOI, loss on ignition at 900°C. NA, not analyzed]

Sample number1 1 2 3 4 5 6

Field number EMP-3-84 EMP-6-84 EMP-7-84 MI-99-5 98STO-309 98STO-310

Major element oxides, in weight percent

SiO2 50.7 51.63 51.26 51.72 51.14 49.54

Al2O3 15.49 14.74 15.81 15.72 15.26 14.85

FeOt 12.91 13.72 10.42 11.76 13.17 13.27

MgO 5.11 5.18 8.31 5.55 6.13 7.05

CaO 8.59 7.56 8.49 10.00 6.85 9.26

Na2O 3.15 3.71 3.71 2.42 3.48 2.75

K2O 1.48 1.15 0.65 1.16 1.42 0.92

TiO2 2.04 1.83 1.06 1.31 1.99 1.87

P2O5 0.22 0.22 0.18 0.17 0.28 0.24

MnO 0.31 0.25 0.11 0.20 0.28 0.26

Total 100 100 100 100 100 100

Volatiles, in weight percent

LOI/H2O+ 1.79 1.84 2.56 1.82 1.7 1.6

H2O– NA NA NA NA 0.1 0.1

S NA NA NA NA 0.03 0.04

Trace elements (ppm)

Sc 36.1 40.1 37.3 38 44 40

V NA NA NA 278 279 217

Co 44 48 48 46 48 54

Cr 58 49 364 81 174 164

Ni 55 58 118 122 68 118

Rb 13.5 26.7 15 31 46.2 31.7

Sr 357 177 310 246 240 330

Cs NA 0.25 0.33 0.30 1.42 1.74

Ba 512 546 200 398 796 338

Y 27 30 15 24 29 26.5

Zr 149 137 72 100 177 133

Nb 19 13 6 9.5 18 15

Hf 3.97 3.78 1.84 2.85 3.51 3.07

Ta 1.59 1.12 0.435 1 1.01 0.939

Th 4.07 5.34 1.5 15 2.72 2.35

U 0.83 1.04 0.25 1.77 0.39 0.45

Zn 96 90 82 69 112 106

Cu 128 122 113 125 192 185

La 24.6 23.8 9.75 24.00 19.6 16.85

Ce 48 45.5 18.6 45.25 38.9 34.25
General Geology 17

Table 1. Whole-rock chemical analyses of Paleoproterozoic mafic rocks from the Emperor Volcanic Complex and diabase near Atkins 

Lake.—Continued

[All measurements in weight percent unless otherwise noted. Major elements recalculated to 100-percent volatile free. Samples 1–4 from Emperor Volcanic 

Complex; samples 5 and 6 from diabase near Atkins Lake. Fe0t is total iron as Fe0. LOI, loss on ignition at 900°C. NA, not analyzed]

Sample number1 1 2 3 4 5 6

Field number EMP-3-84 EMP-6-84 EMP-7-84 MI-99-5 98STO-309 98STO-310

Trace elements (ppm)—Continued

Pr NA NA NA 5.01 NA NA

Nd 25 21.5 9 20.65 20 18.4

Sm 6.13 5.65 2.77 4.91 5.53 4.97

Eu 1.84 1.79 0.94 1.65 1.62 1.51

Gd NA NA NA 4.81 5.37 4.87

Tb 1.011 0.975 0.44 0.860 0.899 0.785

Dy NA NA NA 4.89 NA NA

Ho NA NA NA 0.99 NA NA

Er NA NA NA 2.86 NA NA

Tm NA NA NA 0.401 0.525 0.497

Yb 3.26 3.22 1.97 2.4 3.45 3.04

Lu 0.457 0.493 0.273 0.4 0.494 0.447

Element ratios

[La/Yb]cn2 5.0 4.9 3.3 6.6 3.8 3.7

La/Ta 15.5 21.3 22.4 31.4 19.4 17.9

Th/Ta 2.6 4.8 3.4 20.0 2.7 2.5

Zr/Y 5.5 4.6 4.8 4.2 6.1 5.0

Zr/Nb 7.8 10.5 12.0 10.5 9.8 8.9

Nb3 -31.7 -62.7 -15.6 -188.0 -10.8 -10.8

Ta4 -6.5 -43.9 -11.0 -165.2 -11.8 -6.7

F/F+M5 0.72 0.73 0.56 0.68 0.68 0.65

Mg#6 41.4 40.2 58.7 45.7 45.3 48.6

1

Samples 1, 2, and 3 were analyzed by the U.S. Geological Survey (USGS) using wavelength-dispersive x-ray flrouescence spectroscopy for all ten major 

oxides: J. Tagart, A. Bartel, and D. Siems analysts: energy dispersive x-ray fluorescence spectrometry for Rb, Sr, Ba, Zr, Y, and Cu: J. Jackson, analyst: and 

instrumental neutron activation analysis (INAA) for all other elements including the rare-earth elements: J. Mee, analyst. Sample 4 was analyzed by XRAL 

Laboratories, Don Mills, Ontario using inductively coupled plasma-atomic emission spectroscopy (ICP-MS) for all ten major elements and V, Y, Cu, and Nb: 

all other elements, including the rare-earth elements, were analyzed by the USGS using INAA: C. Palmer, analyst. Samples 5 and 6 were analyzed by Actlabs 

Laboratories, Ltd., Ancaster, Ontario, using ICP-MS and a lithium metaborate/tetraborate fusion method. 

1. EMP-3-84 Massive basalt SW1⁄4SE1⁄4 sec. 20, T.47N., R.43W.

2. EMP-6-84 Diabase NW1⁄4NE1⁄4 sec 26, T.47N., R.44W. 

3. EMP-7-84 Pillow basalt NW1⁄4SE1⁄4 sec. 23, T.47N., R.44W.

4. MI-95-9 Gabbro NE1⁄4NW1⁄4 sec. 26, T.47N., R.44W

5. 98-STO-309 Diabase SW1⁄4NE1⁄4 sec. 20, T.44N., R.5W.

6. 98_STO-310 Diabase SW1⁄4NE1⁄4 sec. 20, T.44N., R.5W.

2

 [La/Yb]cn = La/Yb chondrite-normalized ratio.

3

 Nb = Nb/Chon – (Th/Chon + La/Chon)/2.

4

 Ta = Ta/Chon – (Th/Chom + La/Chon)/2

5

 F/F+M = FeOt/FeOt + MgO.

6

 Mg# = Mg/Mg + total Fe.
18 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

1,000

100

10

1

100

10

1

La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

ROCK/CHONDRITES ROCK/PRIMITIVE MANTLE

Th Nb Ta La Ce Nd Zr Hf Sm Ti Gd Tb Dy Y Ho Er Tm Yb Lu V Sc

A

B

EXPLANATION

Emperor Volcanic Complex

Diabase near Atkins Lake, 

 Wisconsin

Figure 8. Chondrite-normalized rare 

earth element plot (A) and primitivemantle-normalized multi-element plot 

(B) for mafic rocks from the Emperor 

Volcanic Complex and diabase near 

Atkins Lake. Symbols as in figure 7. 

Chondrite normalizing values from 

Nakamura (1974); primitive-mantle 

normalizing values from Kerrich and 

Wyman (1997).

Continental crust is characterized by prominent negative Nb, 

Ta, Ti, and positive Th anomalies (Taylor and McLennan, 

1985); therefore, the assimilation of continental crust during 

the fractional crystallization of the basalt may account for the 

variable inheritance of the crustal compositional anomalies 

(De Paolo, 1981). Ueng and others (1988) showed that crustal 

assimilation and fractional crystallization processes were 

important in the evolution of other Paleoproterozoic basalts 

within the Marquette Range Supergroup in northern Michigan, 

including the Hemlock Formation and Badwater Greenstone. 

Crustal assimilation for the Hemlock basalts is further supported by Nd isotope data (Beck and Murthy, 1991).
Structural Geology 19

Mesoproterozoic Rocks

The Paleoproterozoic strata of the Gogebic iron range 

are overlain by a great thickness of subaerial volcanic and 

sedimentary rocks of the Keweenawan Supergroup, which 

was deposited at about 1.1 Ga in the Midcontinent Rift. This 

report will not discuss these rocks in detail, but will highlight 

some features that are pertinent to the geology of the Gogebic 

iron range. For detailed information on the geology of the 

Keweenawan rocks in the area, see recent maps by Cannon 

and others (1995) and Cannon and others (1996).

The basal unit of the Keweenawan Supergroup in the 

western part of the Gogebic iron range consists of quartzite 

and conglomerate of the Bessemer Quartzite (pl. 1A). The 

Bessemer pinches out eastward and is absent east of Ramsay, Mich., where basalt flows of the Siemens Creek Volcanics overlie rocks of the Marquette Range Supergroup. The 

unconformity between the Marquette Range Supergroup and 

Keweenawan Supergroup marks a hiatus of more than 700 

million years (m.y.). There is only slight angular discordance 

between the two supergroups, however, especially in the central part of the iron range. The discordance indicates that (1) 

little Penokean deformation occurred in the central part of the 

iron range and (2) the present steep northward dips of rocks of 

the Marquette Range Supergroup resulted from Mesoproterozoic deformation rather than from Penokean deformation. In 

the far western part of the iron range, Mesoproterozoic rocks 

were thrust southward along the Atkins Lake-Marenisco fault 

over the Marquette Range Supergroup rocks so that west of 

sec. 14, T. 44 N., R. 5 W., the contact appears to be entirely a 

fault (pl. 1A,B).

Mesoproterozoic intrusive rocks that were intruded at 

approximately 1.1 Ga also have some importance in studies 

of the Paleoproterozoic strata. Gabbro of the Mineral Lake 

intrusion truncates the Paleoproterozoic strata in sec. 14, T. 

44 N., R. 4 W. (pl. 1A). Paleoproterozoic rocks are absent for 

about 9 km to the west but then appear again in sec. 13, T. 44 

N., R. 5 W. and continue for about 20 km farther west before 

being truncated by the Atkins Lake-Marenisco fault. The most 

pervasive effect of the Mesoproterozoic intrusions was the 

formation of a broad contact-metamorphic aureole that altered 

the character of the Ironwood Iron-Formation from about the 

Tyler Forks River to the west. Metamorphism of the Ironwood 

becomes increasingly intense westward. In the Mount Whittlesey area, iron-rich amphiboles are common in more reduced 

parts of the Ironwood. Near the contact with the Mineral Lake 

intrusion southeast of Mineral Lake, the Ironwood contains 

iron-rich orthopyroxenes and local fayalite. The increasing 

magnetite content of the Ironwood west of Tyler Forks River is 

also an expression of increasing metamorphic grade. Rocks in 

the western part of the Gogebic iron range west of the Mineral 

Lake intrusion, have also been substantially metamorphosed. 

The Ironwood is typically an iron-amphibole-garnet-magnetite 

rock in that area. 

Structural Geology

The principal purpose of this section is to describe the 

development of structures in the Paleoproterozoic rocks, 

which vary from large-scale folds and faults to fabric elements 

observed in outcrop. We have divided the Gogebic iron range 

into eastern, central, and western sections, as defined above; in 

this part of the report, we describe the structural development 

of each section. We also compare and contrast the structural 

history of the three sections in order to establish the changes 

that occurred from the very weakly deformed Penokean foreland, characterized by the central part of the range, to the more 

intensely deformed eastern and western sections of the range. 

The structural history of the region began with the intense 

Neoarchean deformation of the basement rocks south of the 

Gogebic iron range at roughly 2.7 Ga, and ended with the 

large-scale northward tilting of the entire region in the closing stages of the Midcontinent rifting at about 1.1 Ga. These 

phases of deformation are discussed only briefly to emphasize 

the aspects that are pertinent to understanding the geology of 

Paleoproterozoic rocks and structures.

A summary of the structural history of the region is given 

in table 2. Alpha-numeric designations used both in the text 

and table 2 are as follows: (1) D designates deformational 

events with the numeric subscripts indicating the order of formation (D1

, D2, and so on); (2) S designates planar structures 

such as bedding (S0), first-generation foliation (S1), secondgeneration foliation (S2

), and so on; (3) L designates linear 

structures, such as first-phase fold axis (or S0/S1 intersection 

lineation) (L1), second-phase fold axis (L2), and stretch lineation (LS

); (4) F designates folds in order of their formation 

(F1, F2, and so on); LA and SA indicate Archean lineations and 

foliations, respectively; and LP and SP indicate Proterozoic 

lineations and foliations, respectively. 

Mesoproterozoic Structures

Mesoproterozoic structures are discussed first in this 

report because their geometric effects on the Paleoproterozoic 

rocks must be removed in order to decipher the geometry 

of the Penokean deformation. Mesoproterozoic strata of the 

Keweenawan Supergroup north of the Gogebic iron range dip 

north at angles ranging from as low as 30o to nearly vertical. 

Although these strata unconformably overlie strata of the Marquette Range Supergroup, they are nearly concordant structurally, the contact being generally a low-angle unconformity. It 

has been clear since at least the work of Schmidt and Hubbard (1972) that the present northward dips of the Marquette 

Range Supergroup strata result from the same Mesoproterozoic monoclinal tilting that affected the Keweenawan strata. 

Schmidt and Hubbard (1972) further proposed that, at the time 

of deposition of the Keweenawan rocks, the Marquette Range 

Supergroup strata dipped gently south because they now gen-
20 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

Table 2. Sequence of tectonic events in the Gogebic iron range.

[D, deformational events with the numeric subscripts indicating the order of formation (D1, D2, and so on); S, planar structures such as bedding (S0), first-generation foliation (S1

), second-generation foliation (S2), and so on; L, linear structures, such as first-phase fold axis (or S0/S1 intersection lineation) (L1), second-phase 

fold axis (L2) and stretch lineation; F, folds in order of their formation (F1, F2, and so on); LA and SA, Archean lineations and foliations, respectively; LP and SP, 

Proterozoic lineations and foliations, respectively]

Event Description Resulting Feature

Mesoproterozoic

Block rotation (D4) Northward rotation of rigid blocks in upper plate 

of listric thrusts

North-facing monocline, small-offset cross faults.

Rifting and volcanism Formation of continental rift basins Midcontinent Rift and thick volcanic and 

sedimentary fill (Keweenawan Supergroup).

Uplift Uplift and peneplanation of Paleoproterozoic strata 

of Penokean fold-and-thrust belt

Unconformity at top of Marquette Range 

Supergroup.

Paleoproterozoic

Block faulting (D3) Reactivation of Presque Isle fault and Little 

Presque Isle fault and related faults. Formation 

of Sunday Lake fault and related faults

Sets of steep faults cutting Midcontinent Rift 

system, spaced shear zones in Archean rocks.

Compressional deformation

D2 Continued thin-skinned coaxial deformation of 

Paleoproterozic strata, including folding of D1

structures. Folding of D1 thrust faults into eastplunging anticline

L2 folds, S2 foliation, Wolf Mountain anticline.

D1 Detachment of Paleoproterozoic strata from 

Archean basement along north-verging thrusts, 

folding of bedding

L1 fold axes, S1 foliation, detachment and thrust 

faults.

Flexural subsidence Deposition of turbidites in advance of northwardmigrating fold-and-thrust beltThick sequence of turbidites of Tyler Formation.

Passage of foreland bulge Uplift and erosion of Menominee Group rocks Unconformity at top of Menominee Group.

Extensional deformation Formation of ancestral Presque Isle trough, 

deposition of iron-formation and volcanic rocks 

in subsiding trough

Presque Isle trough and thick sequence of 

Ironwood Iron-Formation and Emperor Volcanic 

Complex.

Flexural subsidence Formation of shallow-marine stable platform, 

deposition of argillite and quartzite

Transgressive argillite to quartzite sequence of 

Palms Formation.

Uplift and erosion Deep weathering and erosion of Chocolay Group 

and Archean basement rocks

Unconformity at base of Menominee Group.

Subsidence of platform Shallow-marine sedimentation on platform Quartzites and carbonates of Chocolay Group.

Uplift and erosion Peneplanation of Archean fold belt Profound unconformity at base of 

Paleoproterozoic.

Neoarchean

Regional deformation Regional-scale nappes, major faults Thick sequences of partly overturned volcanic 

rocks.

Island-arc formation Eruption of volcanic sequences and emplacement 

of granitic plutons

Volcanic rocks of Ramsay Formation and 

emplacement of Puritan batholith.
Structural Geology 21

erally dip more gently north than do the Keweenawan rocks. 

Restoring the basal Keweenawan strata to the horizontal by 

rotation to the south rotates the Marquette Range Supergroup 

strata to a south dip. Figure 9, modified slightly from Schmidt 

and Hubbard (1972), shows these relations. 

The term “Montreal River monocline” was introduced by 

Cannon and others (1993) for the remarkably thick succession of steeply north-dipping rocks that crops out from the 

Lake Superior shoreline near the mouth of the Montreal River 

at the Wisconsin-Michigan border, southward for nearly 40 

km to the Atkins Lake-Marenisco fault. The monocline has 

exposed a section of rocks at least 30 km thick containing no 

apparent structural repetition. The exposed strata include all 

of the Keweenawan Supergroup units, all of the Marquette 

Range Supergroup units in the Gogebic iron range, and about 

20 km of Neoarchean basement rocks. Cannon and others 

(1993) used structural data based on the present orientation of 

originally vertical diabase dikes and geochronologic evidence 

to date and map the extent of uplift accompanying the rotation. 

They proposed that the Montreal River monocline, including 

all of the Gogebic iron range, is the rotated upper plate of a 

listric thrust fault of near-crustal dimensions whose surface 

trace is the Atkins Lake-Marenisco fault. The fault formed at 

about 1.06 Ga as indicated by Rb-Sr cooling ages of biotite 

from rocks in the rotated and uplifted upper plate (Cannon 

and others, 1993), and is roughly coincident with other major 

thrust faults responsible for partially inverting the deep central 

graben of the Midcontinent Rift.

Penetrative deformation apparently did not occur during this Mesoproterozoic thrusting and rotation. Deformation 

appears to have been limited solely to the rotation of rigid 

blocks. Faults that have small offsets nearly perpendicular to 

the bedding of the Marquette Range Supergroup rocks were 

mapped along much of the Gogebic iron range (for instance, 

Cannon and others, 1995, 1996). Many of these faults can 

be projected northward to align with faults in the lower units 

of the Keweenawan Supergroup. The offset of units ranges 

from negligible to a few hundred meters. The faults are both 

left lateral and right lateral. These faults most likely formed 

during the major Mesoproterozoic thrust faulting and rotation, 

possibly to accommodate changes in the shape of the rigid 

upper plate of the listric thrust. Also, the contact between the 

Mesoproterozoic and Paleoproterozoic strata has been offset 

by these faults (pl. 1B), which indicates that the Little Presque 

Isle and parallel faults in the Presque Isle trough were reactivated at this time.

Atkins Lake-Marenisco Fault

The Atkins Lake-Marenisco fault is one of two major 

faults identified by Cannon and others (1993) as Mesoproterozoic thrust faults, the other being the Pelton Creek fault south 

of the study area. The fault, as defined here, is a composite 

of two long-recognized faults in different parts of the area, 

which, until recent field work, had not been correlated because 

much of their extent is in areas of very limited bedrock exposure where their character and location were poorly known. 

The Marenisco fault was named by Sims (1992) in northern 

Michigan near the eastern end of the Gogebic iron range. The 

Marenisco fault juxtaposed Neoarchean granitic rocks on the 

north with Paleoproterozoic metasedimentary rocks of the 

Marquette Range Supergroup on the south, and was interpreted originally to be a north-verging Penokean thrust fault. 

There are no exposures of bedrock in or near the fault trace 

(pl. 1); its location and nature were interpreted largely from 

aeromagnetic patterns. 

A fault contact between Mesoproterozoic volcanic and 

intrusive rocks on the north and Paleoproterozoic and Neoarchean rocks on the south had been known for many years. 

This fault was named the “Crystal Lake-Atkins Lake fault” 

by Aldrich (1929), who recognized that it was the structural 

equivalent and possibly the direct westward extension of the 

Keweenaw fault in northern Michigan. The name is simplified 

here to Atkins Lake fault because the only exposures of the 

fault are north and east of Atkins Lake in northern Wisconsin (pl. 1A,B). Aldrich (1929) correctly surmised that it is a 

major thrust fault along which Mesoproterozoic volcanic and 

intrusive rocks were thrust southward over Paleoproterozoic 

strata of the Marquette Range Supergroup. Mapping for this 

study shows that the Atkins Lake fault is a westward extension 

of the Marenisco fault, although it is at a different stratigraphic 

position than the Marenisco fault to the east.

The Atkins Lake-Marenisco fault zone is well exposed in 

the gorge of the Marengo River in secs. 14 and 15, T. 44 N., 

R. 5 W. (pl. 1A). The fault zone is best exposed from the 

prominent waterfalls in SW1⁄4 sec. 14 for about 500 m along 

the river to the west. In this area, bedrock bluffs along the 

south side of the river consist of Neoarchean granitic rocks 

and the Paleoproterozoic Bad River Dolomite. There is no 

trace of fault-related deformation detectable in the outcrops. 

Along the north side of the river, exposures of a distinctive fault rock, apparently a finely crushed Neoarchean 

granite, extend from the river bank for several tens of meters 

to the north. This rock appears massive in outcrop, is dark 

gray on fresh surfaces and weathers to a pinkish-gray color. 

Thin sections show that it is composed of a mixture of very 

angular fragments of quartz and plagioclase in a finely comminuted matrix of quartz, plagioclase, chlorite and sericite 

(fig. 10). Chlorite is ubiquitous and has a prominent preferred 

orientation; however, because chlorite composes only a small 

percentage of the rock, it is not abundant enough to impart a 

visible fabric except by microscopic examination (fig. 10). 

The rock is overlain on the north by strongly metamorphosed 

basalt of the Mesoproterozoic Siemens Creek Volcanics. Additional exposures of the rock occur sporadically along the north 

side of the river valley for about 1 km farther west to near the 

center of sec. 15. 

Strongly sheared Archean granite in fault contact with 

the Ironwood Iron-Formation can also be seen sporadically in 

outcrops for a few hundred meters east of the waterfalls. Small 

outcrops of highly sheared metabasalt of the Siemens Creek 

Volcanics may be seen in NW1⁄4 sec. 20, T. 44 N., 
22 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

Figure 9. Block diagram 

illustrating the geometric 

relationship between 

Neoarchean, Paleoproterozoic 

and Mesoproterozoic rocks in 

the central part of the Gogebic 

iron range. Simplified from 

Schmidt and Hubbard (1972).

Bessemer Quartzite

Tyler Formation

Palms

Formation

Neoarchean granite

Fault

Fault

Fault

Stream

Marquette Range Supergroup

(Paleoproterozoic)

Keweenawan Supergroup

(Mesoproterozoic)

Siemens Creek Volcanics

Bessemer Quartzite

Tyler Formation

Ironwood

Iron-Formation

Palms

Formation

Neoarchean granite

fault

fault

fault

Stream

North

R. 5 W. (plate 1A), immediately north of Atkins Lake, and are 

interpreted to mark the Atkins Lake fault. These are the only 

exposures of the Atkins Lake-Marenisco fault along its entire 

surface trace of about 150 km.

The present interpretation of the Atkins Lake-Marenisco 

fault is that it is a Mesoproterozoic thrust fault, as proposed 

by Cannon and others (1993), but its geometry in the western 

part of the Gogebic range is substantially different than shown 

in that report. Recent mapping for the present study suggests 

that the thrust fault, which was deep within Neoarchean rocks 

in the central part of the area, passes upward along a lateral 

ramp to truncate the Mineral Lake intrusion (part of the Mellen Intrusive Complex) on the west and eventually cuts upward 

through the Paleoproterozoic section of the Marquette Range 

Supergroup (pl. 1A,B). In the western part of the area, this 

fault bounds the south side of Mesoproterozoic rocks and has 

emplaced them southward over Paleoproterozoic rocks.

Paleoproterozoic Structures

Structures of Paleoproterozoic age include extensional 

features that formed at the same time as the Marquette Range 

Supergroup sedimentation, and contractional features that 

formed during the ensuing Penokean orogeny. The following 

sections present (1) details of the Paleoproterozoic structural 

evolution, as observed in some key areas of the Gogebic iron 

range, and (2) a study of the contrast between the central portion of the range (described first), where Penokean deformation was minimal, and the more intensely deformed eastern 

and western parts of the range. For each area, the present 

orientation of structural elements is described followed by 

an interpretation of the approximate geometry of the structures prior to Mesoproterozoic rotation which was derived by 

restoring the local basal Mesoproterozoic strata to a horizontal 

orientation. For the western Gogebic iron range, mostly new 
Structural Geology 23

Figure 10. Photomicrographs of rocks from the Atkins LakeMarenisco fault zone on the north side of the Marengo River, 

SW1⁄4, Sec. 14, T. 44 N., R. 5 W. A, Angular to subrounded grains 

of quartz and plagioclase in very fine-grained matrix of chlorite, 

sericite, quartz, and plagioclase. Plane polarized light. B, Material 

similar to A under crossed nichols.

0.5 millimeter

1 millimeter

A

B

observations made during the current investigation are presented; these descriptions also include details of stratigraphy 

and lithology.

Central Gogebic Iron Range

The central part of the Gogebic iron range extends from 

the Sunday Lake fault near Wakefield, Mich. to the Tyler 

Forks River, west of Upson, Wis. (pl. 1B). This part of the iron 

range accounted for a great majority of the previous iron-ore 

production of the range and was mined for nearly its entire 

40-km strike length. Studies conducted during the height of 

mining activity, such as Aldrich (1929), Hotchkiss (1919), 

and Allen and Barrett (1915), provided many details of the 

geology of the central part of the district. Hotchkiss (1919) 

based his studies largely on observations in the extensive 

underground mine workings and, therefore, provided the 

best documentation of the structures of the central part of the 

range where outcrops are rare. Much of the following is based 

on descriptions found in Hotchkiss (1919), along with more 

recent observations, only a few of which were the result of the 

present study.

The central Gogebic iron range, commonly referred to in 

older reports as the “central monocline,” consists of a steeply 

northward-dipping succession of Paleoproterozoic strata that 

are virtually devoid of internal folds and only rarely exhibit a 

penetrative deformational fabric. Paleoproterozoic structures 

are principally faults of four types, as defined by Hotchkiss 

(1919). Their definitions are paraphrased as follows: 

1. Transverse faults—Faults striking nearly perpendicular to 

a rock formation and nearly vertical in dip. 

2. “Eureka”-type faults—Faults striking nearly parallel to 

the strike of a rock formation, or parallel to eastwardpitching dikes, and nearly perpendicular to the beds. 

3. Reverse faults—The Sunday Lake fault is the only one of 

its type.

4. Bedding faults—Faults parallel to the beds of the Ironwood Iron-Formation. 

Three of these types cannot be recognized from outcrop information because of the relatively small displacements and very 

sparse bedrock exposures, but they were well documented in 

the mines and were important in localizing iron ores.

Figure 11, simplified from figure 21 of Hotchkiss (1919), 

shows the geometric relationships of fault types 1, 2, and 4, 

above. Because faults of similar small offsets and of the same 

trend as type 1 faults cut the lower units of the Keweenawan 

Supergroup along the north edge of the Gogebic iron range, 

it is possible that type 1 faults are Mesoproterozoic structures 

and were not formed during the Penokean deformation. They 

do, however, clearly offset bedding faults, which we interpret 

as Penokean compressional features; therefore, if they are 

Penokean structures, then they formed late in the deformational history. Type 2 faults offset bedding faults and are in 

turn offset by transverse faults; therefore, they are intermediate in age between faults of definite Penokean age and faults 

of possible Mesoproterozoic age. The tectonic setting of the 

type 2 faults is not clearly established. When Mesoproterozoic 

tilting is restored by removing the present dip of Mesoproterozoic strata, type 2 faults are nearly vertical. 

Bedding faults (type 4), which have displacements of at 

least 300 m, were well known in mine workings and were best 

described by Hotchkiss (1919). Bedding faults were identified in at least four stratigraphic horizons in the Ironwood 

Iron-Formation, but the principal fault of this type is confined 

to the Yale Member of the Ironwood and was identified over 

the entire strike length of the central Gogebic iron range. In 
24 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

places, this fault constitutes a single surface, but more commonly the displacement was distributed between several 

surfaces, all within the Yale Member. The faults have offset 

the bedding-normal diabase dikes, which were used as markers to determine displacement. Some displacements were 

summarized for various mines by Hotchkiss (1919). Offsets 

everywhere have a sense of movement on north side of the 

fault to the east. Because the offsets are determined for planar 

features, there is no unique solution for total displacement. 

Total displacement may be much larger than the values stated 

by Hotchkiss (1919). Unfortunately, there are no recorded 

observations of kinematic indicators in the fault zones. When 

Mesoproterozoic tilting is restored, these bedding faults dip 

gently south and were apparently low-angle thrust faults. If 

these faults are Penokean compressional features, then they are 

most likely north-directed thrust faults.

The total number of bedding faults in parts of the stratigraphic section that are not exposed in mine workings cannot 

be determined. The relatively small, bedding-parallel displacements, coupled with the very sparse outcrops in the region, 

make these faults nearly impossible to detect except in manmade exposures. They may be common and could account, in 

total, for a considerable amount of Penokean tectonic transport. 

Some evidence indicates contraction and tectonic shortening within the Tyler Formation, although megascopic folds 

are apparently absent. Several outcrops of thin-bedded shale 

and sandstone of the Tyler Formation are located in NE1⁄4 sec. 

28 and NW1⁄4 sec. 29, T. 46. N., R. 2 E. near Hurley, Wis. 

Bedding (S0) in these outcrops is oriented N. 61o E., 80o N. 

(stereoplot O on pl. 1B). Crossbeds in the sandstone show that 

the steeply northwest-dipping beds face toward the northwest. 

S1 cleavage is oriented N. 65° E., 60o N., which is nearly 

parallel to the strike of bedding, but less steeply dipping (S1

on stereoplot E, pl. 1B). Cleavage-bedding intersections trend 

generally N. 65o E. as shown by the diamonds (L1) on stereoplot N (pl. 1B). Rare kink bands deform both cleavage and 

bedding, indicating that Tyler rocks here were affected by two 

phases of Penokean deformation. Basal Keweenawan strata 

only about 100 m to the north dip about 80o N. Restoring the 

Tyler Formation to its pre-Mesoproterozoic orientation indicates that bedding was horizontal and cleavage dipped about 

20o S. There is no evidence of folds within the Tyler Formation in the central Gogebic iron range; all of the reported top 

determinations indicate that north-facing strata and dips are 

similar throughout the unit. The bedding-cleavage intersections are unusual in that they are not parallel to mappable fold 

axes. The well-developed cleavage in these outcrops might be 

caused by tectonic compression during north-directed thrusting on concealed bedding faults, along which upper plates successively overrode lower plates and thus caused successively 

greater shortening upsection. These outcrops with unusually 

well-developed cleavage are within the highest stratigraphic 

levels exposed in the Tyler Formation, which suggests that 

penetrative deformational features may become more intensely 

developed at higher structural levels. On a somewhat broader 

scale, bedding and cleavage in the Tyler Formation between 

Hurley and Upson have a similar orientation, as illustrated in 

stereoplots C and D, respectively, on plate 1B.

In summary, the central Gogebic iron range was subjected 

to only minor deformation during the Penokean orogeny. The 

rocks of the Marquette Range Supergroup were apparently 

tilted as much as 20o S. Faulting occurred in two or three 

phases. First, bedding-parallel thrust faulting produced a small 

northward transport of the upper plates of the thrust faults. 

Later, steep faults cut these thrust faults. No folds are known 

in the central Gogebic iron range, except within a few meters 

of faults, where drag folds having amplitudes and wavelengths 

of a meter or less have been reported. In spite of the apparent 

lack of folds, some rocks high in the stratigraphic section have 

a well-developed slatey cleavage that is apparently related to 

movements during thrust faulting.

Eastern Gogebic Iron Range

The eastern Gogebic iron range extends from the Sunday 

Lake fault near Wakefield, Mich., eastward for about 35 km 

to the western shore of Lake Gogebic (pl. 1B). This section 

contains the most intensely and complexly deformed rocks in 

the Gogebic iron range. The descriptions and interpretations 

presented here result from several studies by others during the 

past several decades, as well as our own studies in the 1990s, 

part of which were presented in Klasner and others (1998). 

Unlike other parts of the Gogebic iron range, the results of 

deformation here can be divided into structures that formed 

during both extensional and contractional phases.

Extensional Phase

A prominent graben, the Little Presque Isle trough 

(Klasner and others, 1998), forms a triangular area in the 

central portion of the eastern Gogebic iron range (fig. 3). This 

graben, bounded by the Presque Isle fault on the southeast 

and the Little Presque Isle fault on the west, was subsiding 

during the deposition of the Ironwood Iron-Formation and the 

eruption of the Emperor Volcanic Complex; thus, it controlled 

the distribution and thickness of both of those units. Because 

of the Mesoproterozoic rotation of the area, the map pattern 

provides an oblique cross section of the now-deformed graben 

and the rocks that filled it. The principal evidence that both the 

Presque Isle and Little Presque Isle faults were active during 

sedimentation is (1) the drastic change in thickness of most 

units of the Menominee Group across the Little Presque Isle 

fault, and (2) the apparent termination of units against the 

Presque Isle fault.

The oldest unit in the graben, the Paleoproterozoic Palms 

Formation, has a poorly constrained thickness that is based on 

only a few drill holes within and outside of the graben. The 

thickness of the Palms does not seem to be strongly influenced 

by the bounding faults, so the graben may not have been active 

at the time of Palms deposition. The thickness of the Ironwood 

Iron-Formation likewise is difficult to determine with preci-
Structural Geology 25

sion because of structural complications, including stratigraphic repetitions by faulting; however, the Ironwood has an 

approximate thickness of 1,000 m within the graben and 650 

m outside of the graben, west of the Little Presque Isle fault. 

Both estimates were made from cross sections by Klasner and 

others (1998) and Prinz (1967), each of whom constructed 

them based on outcrops and drill hole records. 

The most dramatic change in thickness is exhibited by the 

Emperor Volcanic Complex. Immediately west of the Little 

Presque Isle fault, outside the graben, the Emperor is approximately 300 m thick; however, immediately east of the fault, 

within the graben, it is about 1,500 m thick (Klasner and others, 1998). Thus, within the graben the Menominee Group is 

about 1.5 km thicker than immediately outside of the graben. 

This marked change in thickness resulted from subsidence 

of the graben during the deposition of the Ironwood IronFormation and the Emperor Volcanic Complex. Later contractional deformation (discussed below) modified the graben 

substantially, but the relationships discussed above provide 

evidence for its formation during the extensional deformation 

that was coincident with Menominee Group sedimentation. 

In that regard, the graben formed in a manner similar to other 

long-known syndepositional structures in the region such as 

the Marquette trough (Larue and Sloss, 1980).

Contractional Phase

At least two phases of thin-skinned deformation (D1 and 

D2) occurred within the contractional regime in the eastern 

Gogebic iron range. Both were thrusting and folding events 

in which Paleoproterozoic strata were detached either from 

underlying strata or from Neoarchean basement rocks. Slices 

of Neoarchean basement rocks were also incorporated into 

some thrust slices. In addition, later reverse faults during a 

later deformation event (D3) involved both Paleoproterozoic 

and Neoarchean rocks. The Sunday Lake fault (type 3 of 

Hotchkiss, 1919) is the most prominent of these.

Thin-skinned (D1) deformation in the monocline west 

of Little Presque Isle fault resulted in the folding of bedding 

(S0) in Paleoproterozoic strata. The L1 fold axes plunge gently 

east (stereoplot F on pl. 1B). An axial-planar foliation (S1) 

was also formed during this phase (stereoplot G on pl. 1B). 

The results of D1 deformation can be seen in the westernmost 

outcrops of penetratively deformed Paleoproterozoic rocks 

on a hill, locally known as “Radio Tower Hill,” in Wakefield, 

Mich. (SW1⁄4 SE1⁄4, sec. 10, T. 47 N., R. 45 W.). This hill is 

about 2 km east of the Sunday Lake fault. Scour channels 

and crossbeds in thinly bedded argillaceous siltstone of the 

Palms Formation indicate that stratigraphic tops are toward 

the north. Bedding (S0) was folded about gently east-plunging 

fold axes (L1 on stereoplot E, pl. 1B). Axial-planar S1 cleavage is oriented N. 57oW., 35o NE. No D2

 deformation elements 

were observed at this location, but they become prominent in 

outcrops farther east. 

D2 folds were not observed either in rocks of the underlying Ramsay Formation or in the batholith of Puritan Quartz 

Monzonite, both of which constitute the Neoarchean basement 

in this part of the Gogebic iron range. A stereoplot of layering 

in the Ramsay (stereoplot M on pl. 1B) shows that it consistently strikes northwest and dips southwest, which indicates 

that the Ramsay was not folded during the Penokean events. 

Likewise, foliation within the Puritan batholith strikes eastnortheast and dips steeply (stereoplot L on pl. 1B), in contrast 

to trends of Penokean deformation in the overlying strata.

The fault that detached much of the Paleoproterozoic 

strata from Neoarchean basement is exposed in a large outcrop 

in the SE1⁄4 sec. 18, T. 47 N., R. 44 W. (Klasner and others, 

1998). The unconformity between the Ramsay Formation and 

Tyler Formation

Ironwood

Iron- Formation

Palms

Formation

“Eureka”-type fault

bedding fault

transverse fault

North

Figure 11. Schematic block diagram 

illustrating the geometry of fault sets in 

the central Gogebic iron range. Simplified 

from Hotchkiss (1919). Black units are 

diabase dikes.
26 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

the Sunday Quartzite can be seen in this outcrop. Crossbeds in 

the Sunday Quartzite indicate that it faces north. A mylonitic 

shear zone parallel to bedding is present near the top of the 

Sunday Quartzite. Structural vergence in the shear zone was to 

the north, which suggests that the Paleoproterozoic rocks were 

detached from the underlying Neoarchean rocks (and the basal 

part of the Sunday in this location) on this originally gently 

dipping shear zone. 

D2 deformation west of the Little Presque Isle fault 

resulted in the folding of S1 axial-planar cleavage about L2 fold 

axes that are coaxial with the D1 fold axis (stereoplot G on 

pl. 1B). Although D2 structures formed after D1 folding, they 

do not necessarily record a discrete event; instead, D1 and D2

structures likely were formed in a continuum of north-verging, 

convergent deformation. 

The same sequence of deformation is present in rocks 

east of the Little Presque Isle fault, within the Presque Isle 

trough. Stereoplot H (pl. 1B) shows that bedding was folded 

about gently east-plunging L1 fold axes in a manner similar to 

that shown in stereoplot F (pl. 1B) west of the Little Presque 

Isle fault. S1 axial-planar foliation related to F1 folds was, in 

turn, folded about L2 fold axes that are coaxial with L1 (stereoplot I on pl. 1B), which is the same as that shown west of the 

fault. The gently plunging axis of the Wolf Mountain anticline 

(fig. 3) is parallel to the L1 and L2 fold axes found both east 

and west of the Little Presque Isle fault. Also, shear zones and 

faults of the D1 phase were folded about the anticline (Klasner and others, 1998), which indicates that the anticline is a 

D2 structure. Such structures are not found in the Neoarchean 

basement rocks, indicating that rocks in the Wolf Mountain 

anticline were detached from Neoarchean basement. 

A final deformation event (D3) in the eastern Gogebic 

iron range resulted in several sets of faults along which Paleoproterozoic and Neoarchean rocks were broken into blocks. 

Some of these faults are certainly of Paleoproterozoic age 

and, therefore, formed during the Penokean orogeny;, many 

of the faults, including the Little Presque Isle fault, caused 

an offset of the basal units of the Keweenawan Supergroup 

(Mesoproterozoic), which clearly indicates that the faults were 

reactivated during the Mesoproterozoic deformation. The 

Little Presque Isle fault produced nearly 500 m of offset of 

the surface trace of the base of the Siemens Creek Volcanics 

(Mesoproterozoic). This Mesoproterozoic fault movement 

may also be responsible for the observed relationships farther 

south along the fault, but Mesoproteozoic deformation is difficult to isolate from earlier deformation in those areas. 

The Sunday Lake fault (pl. 1B) is the most prominent of 

the D3 faults. It can be traced from well within Neoarchean 

basement in the eastern part of the Gogebic iron range, 

through the Marquette Range Supergroup to the west, to a 

point where it is truncated by the unconformity at the base of 

the Mesoproterozoic rocks near Wakefield, Mich. The fault 

resulted in about 2 km of left-lateral offset of surface traces 

of Paleoproterozoic rocks, but did not offset the basal units 

of the Siemens Creek Volcanics. The Sunday Lake fault was 

apparently a major fault that formed late during the Penokean 

orogeny. A few outcrops of Neoarchean rocks along the fault 

trace have a vertical to steeply south-dipping foliation. Studies 

of the fault during mining indicated that it is nearly vertical (Hotchkiss, 1919). If 65o of Mesoproterozoic rotation is 

removed, the Penokean orientation of the fault is a north-striking and moderately east-dipping fault plane in which the upper 

plate had a significant westward reverse component of displacement, but for which total displacement is unconstrained. 

Elsewhere in the Penokean orogen, a deformational sequence 

that began with thin-skinned deformation and culminated in 

basement-block uplift along reverse faults has been proposed 

(Cannon, 1973; Klasner, 1978). Faults that bound basement 

blocks have diverse orientations and are commonly highly 

divergent from trends of older contractional structures. The 

Sunday Lake fault is, therefore, one of that family of late-tectonic reverse faults and the deformational sequence observed 

in the eastern Gogebic iron range is similar to that established 

elsewhere in northern Michigan. The Mesoproterozoic rotation 

of the eastern Gogebic iron range has provided a unique crosssectional view of structures that can be seen elsewhere only in 

plan view.

Restoration of the Presque Isle trough to its preKeweenawan orientation provides a view of the trough, its 

bounding faults, and its contained structures from a different 

perspective than the current cross-sectional view seen in figure 

3. To restore the trough to its pre-Keweenawan orientation, the 

structures were rotated 50o S. around an east-oriented axis (fig. 

12). In this process, the fold axes’ orientations were changed 

only slightly because they were nearly parallel to the axis of 

rotation. Likewise, the strike of the Presque Isle fault changed 

only slightly because it is nearly parallel to the rotation axis. 

The fault’s dip, however, changed from 60o SE. to about 80o 

NW. (compare stereoplot A and stereoplot B on fig. 12). In 

the same manner, the strike of the gently southeast-dipping 

Little Presque Isle fault changed from N. 20o E. to N. 65o E. 

and its southeast dip steepened to 60o(compare stereoplot C to 

stereoplot D on fig. 12); thus, when the Mesoproterozoic rotation was removed, the Little Presque Isle fault and the Presque 

Isle fault diverge only slightly in strike and define a gently 

east-northeast-plunging trough bounded by steep faults (see 

E on fig. 12) which is similar to the west-plunging Marquette 

trough to the east. 

Mesoproterozoic rotation of the trough has some other 

implications. The Little Presque Isle fault is one of a family of 

faults that crosscut the Presque Isle trough (fig. 3). The faults 

extend northward, offsetting the unconformity at the base of 

the Mesoproterozoic rocks, indicating that the Little Presque 

Isle fault set was reactivated during the Mesoproterozoic 

deformation; therefore, the Mesoproterozoic reactivation of 

these faults probably occurred during south-directed thrusting 

on the Atkins Lake-Marenisco fault. The steep faults may have 

formed in order to accommodate the changing shape of the 

upper plate of this regional listric thrust fault.

The easternmost extent of our studies included outcrops 

of the Copps Formation near Lake Gogebic. As shown in plate 

1B and fig. 3, the unconformable base of the Copps Forma-
Structural Geology 27

tion truncated older parts of the Marquette Range Supergroup. Eastward from sec. 23, T. 47 N., R. 43 W., the Copps 

lies directly on Neorchean granitic rocks. Rocks in Gogebic 

County Park near the south end of Lake Gogebic (NW1⁄4,SW1⁄4

sec. 3, T. 40 N., R. 42 W.) contain many structural elements which are summarized here for comparison with those 

described in the Gogebic iron range farther to the west. At the 

park, the Copps Formation consists of medium- to fine-grained 

graywacke with graded beds, refracted cleavage, and stretched 

concretions, all of which provide information on the structural 

history of the easternmost part of the Gogebic iron range. 

The initial phase of deformation resulted in F1 folds that 

plunge 50o N., 75o E. within the Copps. As shown in stereoplot J of plate 1B, poles to bedding (S0

) indicate little evidence of small-scale folding and only two L1

 fold axes were 

observed; however, there is abundant evidence for folding 

of the S1 cleavage that is axial planar to F1 folds. S1 is folded 

about L2 fold axes that plunge 50o at S. 60o E. (stereoplot K 

on pl. 1B), and S2 axial-planar foliation is also folded, both of 

which indicate that there were additional phases of deformation in these rocks. The long axes of the stretched concretions 

within the Copps plunge 59o, S. 24o E; thus, the Copps at the 

easternmost end of the Gogebic iron range was deformed at 

least twice about east-trending fold axes. When corrected for 

Keweenawan rotation, both the L1 and L2 fold axes plunge 

moderately toward the northeast and roughly parallel to the 

L1 and L2 folds in the Presque Isle trough, which suggests that 

both areas were subjected to the same deformational events.

Summary of Eastern Gogebic Iron Range Structure

 The eastern Gogebic iron range was substantially more 

tectonically active during the Penokean orogeny than was the 

central part of the iron range. The eastern Gogebic iron range 

is unique in that it preserves clear evidence of extensional 

deformation that coincided with deposition of the Menominee 

Group. Growth faults bounding a graben (the Presque Isle 

trough) exerted a strong influence on the character and thickness of the Ironwood Iron-Formation and Emperor Volcanic 

Complex. Contractional deformation was also pronounced and 

consisted of an early, thin-skinned phase that resulted in folds 

and cleavage that were geometrically uniform throughout the 

entire eastern part of the iron range, but were more widely and 

intensely developed to the east. The contractional deformation 

was absent in Neoarchean basement rocks; the fabrics of the 

Neoarchean rocks still reflect the Neoarchean deformation, 

thereby attesting to the thin-skinned character of this phase 

of deformation. A final Paleoproterozoic deformational phase 

resulted in reverse faults and the uplift of the fault blocks. The 

Sunday Lake fault is the most prominent of these faults and 

appears to have thrust Neoarchean and Paleoproterozoic rocks 

to the west or northwest.

Figure 12. Diagram illustrating the restoration of the 

Paleoproterozoic geometry of the Presque Isle trough constructed 

by compensating for the Mesoproterozoic northward rotation. 

The Presque Isle trough in Paleoproterozoic time was an eastwest-trending graben bounded by normal faults, the Presque Isle 

and Little Presque Isle faults.

A B

C

E

D

U

D

D

U

North Little Presque Isle fault

Presque Isle fault

Ancestral Presque Isle trough

N 15° E

N 59° E

Pre-restoration Post-restoration

N 80° E

N 80° E

Pre-restoration Post-restoration

Orientation of the Little Presque Isle fault

Orientation of the Presque Isle fault

After the Paleoproterozoic deformation event, the entire 

eastern Gogoebic iron range and the adjacent Neoarchean 

basement were tilted northward as rigid blocks at about 1,100 

Ma (Cannon and others, 1993). This resulted in the present 

map pattern, which may be interpreted as an oblique cross 

section of the Paleoproterozoic structures (such as the Presque 

Isle trough). The uplift also reactivated the Paleoproterozoic 

faults of the Little Presque Isle fault set.
28 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

Western Gogebic Iron Range

The western Gogebic iron range extends from Tyler 

Forks River for about 55 km westward to the point where the 

Marquette Range Supergroup rocks have been truncated by 

the Atkins Lake-Marenisco fault. The Tyler Forks River was 

chosen as the eastern boundary for the western Gogebic iron 

range because (1) it is approximately the western limit to 

which the classical internal stratigraphy of the central Gogebic 

iron range has been traced with confidence, and (2) it is the 

approximate eastern limit of intense Penokean deformation in 

the Marquette Range Supergroup. 

The western Gogebic iron range differs from the central 

Gogebic iron range both in lithologic character (described 

above) and in structure. Compared to the central and eastern 

parts of the range, the western Gogebic iron range has been 

much less studied, with the exception of unpublished mining 

company studies. The most recent comprehensive description 

was by Aldrich (1929). As a result, we examined the western 

part of the range in some detail from 1990 to 1994, during 

which time we compiled maps at 1:24,000 scale and revised 

some aspects of the previously published geology. This work 

was aided by access to very detailed data gathered in the late 

1950s by United States Steel Corporation (USS) as part of 

a detailed economic evaluation of iron resources in the area 

from Tyler Forks River westward to the Bad River at Penokee 

Gap. That work included 1:2,400-scale geologic mapping, 

ground magnetic surveys, and drilling. Some of that has been 

generalized in order to incorporate it into this report. The new 

1:24,000-scale detailed map (pl. 1A) of the western Gogebic 

iron range forms the basis for the following descriptions of the 

structural history of this part of the range.

The principal new finding of this study is the recognition 

of faults that are here interpreted as thrust faults that formed 

during the Penokean orogeny. These faults are nearly parallel 

to bedding in the Marquette Range Supergroup rocks and lie 

both at the base of the section where they form basal décollements on the Neoarchean basement, and within the lowermost 

few hundred meters of the Marquette Range Supergroup rocks. 

These faults, in places, splay up in the section and result in 

the repetition of units such as is seen on Mount Whittlesey. 

Because of northward rotation of the western Gogebic iron 

range during the Mesoproterozoic, both the Marquette Range 

Supergroup rocks and the faults within them now have steep 

north dips. Folds of outcrop scale and larger are also common 

in parts of the western Gogebic iron range and are especially 

prominent near Mineral Lake, where a combination of upright 

folds formed during the Penokean orogeny and Mesoproterozoic rotation produced the present geometry of recumbent 

folds. The following paragraphs describe these structures from 

the east, near Tyler Forks River, to the west, ending where 

the Marquette Range Supergroup rocks were truncated by the 

Atkins Lake-Marenisco fault.

Immediately west of Tyler Forks River tight chevron 

folds were observed in a few outcrops of the Ironwood IronFormation, which indicates Penokean deformation; Aldrich 

(1919) described tightly folded rocks in this same area. Most 

folds have axial planes inclined moderately to the south. The 

Marquette Range Supergroup rocks appear to be autochthonous over Neoarchean basement. Outcrops of the lower part 

of the Palms Formation, which is probably only a few meters 

above the basal contact with Neoarchean basement, show no 

tectonic fabric. Within a few hundred meters east of the map 

area of plate 1A, a well-exposed unconformity was described 

by Aldrich (1929) in which tens of meters of paleotopographic 

relief was observed on the upper surface of the Neoarchean 

rocks. The basal beds of the Palms Formation fill the topographic depressions on that surface. The basal décollement 

appears to be absent this far east in spite of its clear presence 

only a few kilometers to the west, as described below. 

We have inferred that a basal décollement extends 

eastward from the Mount Whittlesey area, where it is well 

exposed, to approximately NE1⁄4 sec. 11, T. 44 N., R. 2 W., a 

few hundred meters east of its easternmost exposure. The eastern extent of the detachment is not well constrained by outcrop 

information because there are no additional exposures of the 

basal contact of the Marquette Range Supergroup for several 

kilometers to the east. Westward from Ballou Creek, there are 

abundant outcrops in the Mount Whittlesey area. The detailed 

studies by USS guided our study of that area and were reinterpreted in several areas in order to produce the map on plate 

1A. The geology in this area is of sufficient complexity that, 

even with very detailed maps, magnetic surveys, and drilling 

by USS, and the recent work for this report, the structure is not 

unequivocally delineated.

Mount Whittlesey is the easternmost intensely deformed 

area of the western Gogebic iron range and thus provides an 

example of the nature of the leading edge of the Penokean 

foreland deformation belt. The area presents a structural pattern in which the Marquette Range Supergroup rocks were 

detached from Neoarchean basement. Thrust faults splayed up 

from the basal décollement and produced structural repetitions 

of the Ironwood Iron-Formation. Folds of various scales are 

also present. The map pattern yields an oblique cross section 

of the Penokean structures which approximately parallels 

the trend of those structures. Cross section B–B′ on plate 1A

shows a transverse view of the structures, both as presently 

configured (section A) and with Mesoproterozoic rotation 

removed to approximate the original Penokean geometry (section B).

The basal décollement is best seen in exposures in the 

NW1⁄4 sec. 11, T. 44 N., R. 2 W., where the basal breccia unit 

of the Palms Formation overlies dacite breccia of the Ramsay 

Formation. The dacite breccia has a strong tectonic fabric that 

resulted from Neoarchean deformation. Breccia fragments 

were highly stretched and define a lineation plunging 20o to 

30o NW. A weak to moderate foliation strikes generally WNW. 

and dips about 50o S. When corrected for Mesoproterozoic 

rotation, the dips of the lineations and foliation become steeper 

and, therefore, consistent with structures in other Neoarchean 

rocks in the region. Near their contacts with the Marquette 

Range Supergroup, a prominent shear fabric is present in both 
Structural Geology 29

the Neoarchean volcanic rocks and in the Palms Formation. 

This foliation is in the lowermost few meters of the Palms, but 

extends for 100 m or more into the Ramsay Formation. The 

foliation strikes about N. 80o E., dips from 20o to 75o N., and, 

in some outcrops, overprints south-dipping Neoarchean foliation. One additional small outcrop in the NW1⁄4 sec. 16, T. 44 

N., R. 2 W. shows a basal shear zone in which the Bad River 

Dolomite, a unit too small to show on plate 1A, has highly 

sheared and silicified rocks near its base; the shearing and 

silicification presumably formed during thrusting on a basal 

detachment zone.

Several additional thrust faults account for much of 

the structural complexity of the Mount Whittlesey area. The 

thrusting appears to have cut Neoarchean rocks, at least 

locally, such as in SE1⁄4 sec. 10, T. 45 N., R. 2 W., where a 

sliver of Palms Formation and Ironwood Iron-Formation is 

structurally overlain by volcanic rocks of the Ramsay Formation. A thrust fault that splayed from the basal décollement 

on the south side of Eagles Peak migrated up section to the 

east and repeated the Ironwood Iron-Formation; this repetition 

accounts for the unusual apparent thickness of the Ironwood 

on Mount Whittlesey. Still farther east, that thrust fault again 

became the basal thrust. The existence of these faults was 

documented best by USS mapping that showed offsets of units 

within the Ironwood Iron-Formation. These internal units are 

not shown on plate 1A.

Folds are common in the Ironwood Iron-Formation on 

Mount Whittlesey. For instance, the hill containing the eastern 

part of the lowermost thrust panel in the SW1⁄4 sec. 10, T. 44 

N., R. 2 W., is underlain by Ironwood that has been folded into 

a west-plunging syncline. Additional folds that have wavelengths and amplitudes of roughly 10 m are well exposed in an 

abandoned railroad cut at the Berkshire mine in the SE1⁄4SW1⁄4

sec. 9, T. 44 N., R. 2 W. The detailed maps by USS indicate 

several areas where large-scale folds were inferred within individual lithologic units of the Ironwood, but the reason for this 

interpretation is not obvious from the information examined 

for this report; therefore, the folds are not shown on plate 1A. 

Folds at outcrop scale are also widely distributed. Stereoplot Q 

(pl. 1B) of poles to bedding (S0) shows that beds were folded 

about L1 axes that plunge gently westward. Axial-planar S1

foliation was, in turn, folded about L2 fold axes that plunge

gently at S. 72o W., as shown on stereoplot P (pl. 1B).

Restoring the structures to their Paleoproterozoic orientation by removing 70o of Mesoproterozoic rotation produced 

relatively little change in the fold-axis orientations because 

they plunge roughly parallel to the axis of rotation. Corrected 

L1 fold axes plunge gently toward the east, whereas L2 axes 

plunge gently southwest. At the eastern end of Mount Whittlesey, the basal shear zone strikes about N. 70o E. and dips 

60oto 75o NW., and bedding in the overlying Paleoproterozoic 

rocks dip 45oto 67o NW. When corrected for Keweenawan 

rotation, the shear zone becomes nearly horizontal and bedding in the overlying Paleoproterozoic rocks dips gently 

southward. 

West from Mount Whittlesey, the map pattern of the Marquette Range Supergroup rocks shows a north-facing monocline that is similar to the structure of the central Gogebic 

iron range, yet folds and faults are widespread although not 

of sufficient magnitude to have repeated map units. In the 

Penokee Gap area, where the Bad River transects the range 

and for about 5 km farther west, mapping is based on both 

the USS data, which extends westward through sec. 14, and 

our own observations farther west. A description of this area 

was included in a field guide (Klasner and LaBerge, 1996) 

and much of the following is an excerpt from that guide. The 

north-facing monocline of Paleoproterozoic rocks is evident 

in stereoplot B (pl. 1B). The Paleoproterozoic rocks have a 

continuous basal unit of chert breccia in the Palms Formation. 

Small remnants of the Bad River Dolomite occur locally, such 

as along the Bad River in SW1⁄4 sec. 14, T. 44 N., R. 3 W., but 

are too small to show at the scale of plate 1A. The remainder 

of the rocks consist of the typical sequence of Palms Formation overlain by the Ironwood Iron-Formation and Tyler 

Formation. 

Sparse outcrops of Neoarchean rocks south of the range 

have a nearly vertical, east-striking foliation and parallel shear 

zones. Neither of these structural elements was observed in 

the overlying Paleoproterozoic rocks; therefore, they are the 

result of Neoarchean deformation. Two closely related phases 

of Penokean deformation were recognized in the Penokee Gap 

area (D1 and D2). D1 resulted in thin-skinned, north-verging 

thrusting that (1) detached the Paleoproterozoic strata from 

underlying Neoarchean basement rocks, or (2) displaced 

Paleoproterozoic rocks on shear zones approximately parallel to bedding. D1

 deformational features are widespread in 

the Penokee Gap area. They include folds (F1) in bedding 

(S0) about gently northwest-plunging axes in the Ironwood 

Iron-Formation and Palms Formation (stereoplot A on pl. 1B). 

Similar folds as well as axial-plane cleavage are locally present in the Tyler Formation. Fracture cleavage formed locally in 

the Ironwood Iron-Formation (stereoplot B on pl. 1B).

Prominent shear zones shown by S1 mylonitic foliation 

are present in the Palms Formation and at the contact between 

the Ironwood Iron-Formation and Tyler Formation. The Palms 

is a thin-bedded, quartzose argillite immediately above its 

basal breccia member. The unit contains a mylonitic foliation approximately parallel to bedding, which is apparently a 

zone of concentrated shear across which overlying units were 

detached from the basal Paleoproterozoic and Neoarchean 

basement rocks. A similar shear zone lies along the base of 

the Tyler Formation. The shear zone in the Palms Formation 

verges northward, as indicated by the displacement and folding of individual beds across the mylonitic foliation. Northward vergence in the Tyler Formation is also indicated by 

locally observed small-scale folds.

D2 deformation was characterized by the continued 

formation of north-verging thrust faults and the continued 

deformation along the S1 shear zones. Within the Palms 

Formation, prominent F2 kink folds in S1 mylonitic foliation 

plunge gently to the northwest (stereoplot B on pl. 1B) and 
30 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

are prime examples of the D2 deformation. F2 kink folds were 

not observed in the Tyler Formation, but S1 foliation appears 

to have been folded about northwest-plunging axes that are 

mostly parallel to F2 axes (stereoplot B on pl. 1B). 

When corrected for the Mesoproterozoic rotation, the 

Paleoproterozoic rocks dip gently to moderately to the south. 

The shear zones in the Palms and Tyler also dip to the south 

and clearly seem to have originally been north-verging thrust 

faults. Both L1 and L2 folds plunge gently east to slightly 

southeast, roughly parallel to the L1 and L2 folds in the eastern 

Gogebic iron range. 

Still farther west, the area south and east of Mineral 

Lake (pl. 1A) contains the most intensely folded Paleoproterozoic rocks in the western Gogebic iron range. Outcrops of 

the Ironwood Iron-Formation and Palms Formation contain 

many northwest-plunging folds that have wavelengths of a 

few hundred to about five hundred meters. Smaller-scale folds 

are also common at outcrop scale. Fold axes, as measured in 

outcrop-scale folds, bear roughly N. 45o W. and plunge 45o 

NW. Axial surfaces have moderate to shallow southwesterly 

dips. Figure 13, for instance, shows a tight fold in the Ironwood Iron-Formation, which is typical of the general geometry of folds in this area. Axial surfaces are inclined to the 

south and anticlines have gently south-dipping south limbs and 

steep to overturned north limbs. The folds are apparently rootless and the Paleoproterozoic rocks are apparently detached 

from the Neoarchean basement rocks. A basal detachment is 

not exposed here, but some of the rocks in the southernmost 

outcrops of the Palms Formation, such as at the southeast 

corner of sec. 18, T. 44 N., R. 3 W., have a strong foliation 

not generally common in the unit, which suggests that there is 

probably a shear zone immediately south of these outcrops. A 

splay from the basal detachment has thrust Palms Formation 

over the Ironwood Iron-formation in the SE1⁄4 sec. 8, T. 44 N., 

R. 3 W. The only outcrop of Neoarchean rocks mapped in the 

area is in the NW1⁄4 sec. 24, T. 44 N., R. 4 W. Here, the rock is 

a massive granitoid of the Puritan batholith and is only about 

100 m from the probable basal detachment. The lack of a 

Penokean fabric and the proximity of this outcrop to intensely 

folded Marquette Range Supergroup rocks further indicate the 

thin-skinned nature of the deformation. When the effects of 

Mesoproterozoic rotation are removed, the folds are upright, 

the basal detachment is subhorizontal (pl. 1A, cross section 

A–A′), and fold axes have shallow plunges in a more westerly 

direction than their present orientation.

An unusual feature of the Mineral Lake area is a thick 

section of the Bad River Dolomite found in the NW1⁄4 sec. 24, 

T. 44 N., R. 4 W., which appears to be several hundred meters 

thick. The rock consists of cherty, dolomitic marble that is 

typical of the Bad River elsewhere in the Gogebic iron range. 

This section is the only known preserved Bad River for many 

kilometers along the strike of the range and appears to have 

been overridden by the basal detachment of the main part of 

the range. This section of Bad River Dolomite may be a fault 

block that was dropped down early in the history of the area, 

before the Bad River-Palms erosion interval, and thus was 

preserved because of its structural position. If that is the case, 

the Bad River must have been an extensive unit in the area and 

was removed by pre-Palms dissolution and erosion. Tight folds 

are present in some outcrops and have a geometry similar to 

that seen to the north in the Ironwood, which suggests that the 

Bad River was folded during the same events as the overlying 

rocks. 

The Marquette Range Supergroup was truncated on the 

west by the base of the Mineral Lake intrusion (a large mafic 

pluton). Metamorphism of the country rock was intense and, 

within a few hundred meters of the contact, the Ironwood was 

converted to a coarse-grained granoblastic rock composed of 

quartz (recrystallized chert), magnetite, orthopyroxene, and 

local fayalite. Near the intrusive contact, bedding of the Ironwood strikes to the northeast, nearly parallel to the base of the 

intrusion, which suggests that local deformation resulted from 

emplacement of the pluton.

Gabbro of the Mineral Lake intrusion truncated the Marquette Range Supergroup rocks in the NW1⁄4 sec. 14, T. 44 N., 

R. 4 W., and the strata are absent for about 10 km to the west 

where they reappear, first as an enclave within the Mineral 

Lake intrusion, and finally as a continuous section in the footwall of the Atkins Lake-Marenisco fault. 

Some important relationships are shown in outcrops 

along the Marengo River in secs. 14, 15, and 16, T. 44 N., R. 

5 W. The geology of the Marengo River area was discussed 

above in the section on Mesoproterozoic structures because 

of the excellent exposures of the Atkins Lake-Marenisco fault 

in the area. The following paragraphs explain some additional 

aspects of the Paleoproterozoic geology. 

The Marengo River area differs from other parts of the 

Gogebic iron range in being mostly in the footwall of the 

Atkins Lake-Marenisco fault; however, it is also in the hanging wall of the related Pelton Creek fault. Because of this 

relationship, the area apparently was rotated to the northwest 

somewhat independently of the rocks in other parts of the iron 

range. Paleoproterozoic rocks are preserved in two areas. An 

Figure 13. Small scale recumbent fold in Ironwood IronFormation in NE1⁄4, sec. 24, T. 44 N., R. 4 W. View looking west.
Economic Geology 31

enclave of strongly metamorphosed Paleoproterozoic rocks 

consisting of the Ironwood Iron-Formation and unnamed gabbro was mapped within the Mesoproterozoic Mineral Lake 

intrusion. The enclave underlies a triangular area in secs. 13, 

14, 23, and 24, T. 44 N., R. 5 W. and is truncated on the west 

by the Atkins Lake-Marenisco fault. The Ironwood is mostly 

an evenly-bedded cherty, iron-amphibole-rich rock containing variable amounts of magnetite. No deformational features 

were observed in the relatively few outcrops of the iron-formation, all of which contain beds that dip moderately to steeply 

to the north-northwest. 

The Paleoproterozoic rocks south of the Atkins LakeMarenisco fault form a northwest-dipping monocline above 

Neoarchean granitic rocks. The basal unit of the Marquette 

Range Supergroup (the Bad River Dolomite) lies unconformably on the granite. The basal unconformity is well exposed 

near the southeast corner of sec. 15, T. 44 N., R. 5 W. At that 

location, neither the dolomite nor the granite show tectonic 

fabric. The dolomite generally does not show deformation fabric in the widespread outcrops and in a quarry to the 

southwest. These observations indicate that the Bad River 

was essentially unaffected by Penokean deformation and that 

the 45o to 55o dips are a result of Mesoproterozoic rotation. A 

Mesoproterozoic diabase dike near the quarry in sec. 22, T. 44 

N., R. 5 W., dips 50o S., which indicates substantial Mesoproterozoic rotation of this structural block. Although there is 

no unique solution for the amount and direction of rotation, it 

seems likely that the Bad River Dolomite was nearly flat-lying 

prior to the Mesoproterozoic deformation. 

Evidence for substantial Penokean deformation appears 

higher in the stratigraphic section. A broad zone of intense 

shearing is well exposed in the argillite near the base of the 

Palms Formation that is parallel to its contact with an underlying sill of metamorphosed gabbro in the SW1⁄4 sec. 15, T. 

44 N., R. 5 W. (pl. 1A). The stratigraphically lowermost 50 

m of the shear zone largely obliterates the bedding, except 

for the most quartzose of the layers, which are tightly folded. 

Those folds were largely dismembered by attenuation of their 

limbs. The intensity of deformation diminishes upsection, but 

the effects of shearing are seen for at least an additional 100 

m. Small folds indicate a northward vergence. The shearing 

is roughly parallel to bedding in the Palms and is generally 

parallel to the strike of the rock units in the area. The shear 

zone, therefore, appears to be a major detachment surface 

along which the Palms Formation, Ironwood Iron-Formation, 

and metagabbro sills were thrust over undeformed Bad River 

Dolomite, and would have been a low-angle detachment prior 

to Mesoproterozoic tilting. 

The Palms Formation and Ironwood Iron-Formation overlying the shear zone constitute a northwest-facing monocline, 

but show small folds in some outcrops. The few observed fold 

axes plunge at moderate angles to the northwest and restore to 

shallowly plunging folds when the Mesoproterozoic rotation is 

removed. The general trend of Paleoproterozoic units is interrupted by a plug of massive, pink, coarse-grained granite in 

sec. 20, T. 44 N., R. 5 W. The granite does not show deformational fabric and is lithologically identical to granitic phases of 

the Mesoproterozoic Mellen Intrusive Complex; therefore, we 

interpret it to be a Mesoproterozoic intrusion.

The westward continuation and eventual termination of 

the Marquette Range Supergroup west of the Marengo River 

area is poorly defined. The westernmost bedrock exposures 

of the strata are in sec. 26, T. 44 N., R. 6 W. The Ironwood 

Iron-Formation in that area consists of thin-bedded, slightly 

cherty rock composed mostly of iron-amphiboles, and only 

minor amounts of magnetite. Drill holes in the area also 

indicated similar weakly magnetic iron-formation interlayered 

with black shale. Additionally, metagabbro sill crops out in 

sec. 26. We have attempted to trace the Ironwood westward by 

magnetic surveys (including ground traverses) but the anomalies are very weak and indicate that the magnetite content 

continues to decline to the west. Our interpretation is that the 

Marquette Range Supergroup rocks were truncated against 

the Atkins Lake-Marenisco fault in secs. 33 and 34, T. 44 N., 

R. 6 W. The westward extension of the Bad River Dolomite 

is even more problematic. There are no exposures of it west 

of the abandoned quarry in sec. 22, T. 44 N., R. 5 W. so the 

projected continuation of it to the west and its truncation by 

the Atkins Lake-Marenisco fault is speculative. 

Economic Geology

The Gogebic iron range was a major iron ore producer 

for 90 years from 1877 to 1967. During that period, about 

325,091,000 tons of ore were shipped from about 40 mines 

between Upson, Wis., on the west and Wakefield, Mich., on 

the east. This 40-km-long belt was mined nearly continuously, mostly from underground workings. The Gogebic iron 

range ranks third nationally in natural ore production behind 

the Mesabi iron range in Minnesota and the Marquette iron 

range in northern Michigan. Ore bodies are masses of nearly 

pure iron oxides and hydroxides that formed as secondary 

concentrations (so-called “natural ores”) within the Ironwood 

Iron-Formation. The term “natural ore” distinguishes these 

deposits from lower-grade concentrating ores (commonly 

called “taconite”) which must be concentrated in order to 

produce a useable product. Taconite has never been produced 

from the Gogebic iron range, but a detailed assessment of its 

western part from Upson to Mineral Lake indicates that this 

35-km-long belt contains a resource of about 3.7 billion tons 

of material (Marsden, 1978) that was deemed to be economically mineable at the time of that assessment, making the 

western Gogebic iron range one of the largest undeveloped 

taconite resources in the United States. 

Natural Iron Ores

Although the presence of iron-bearing rocks was known 

as early as 1848 from observations by linear surveyors near 

Upson, Wis., iron ore was not discovered until many years 
32 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

later when a trapper discovered ore beneath an overturned tree 

at what became the Colby mine near Bessemer, Mich. (Reed, 

1991). The first mining commenced in 1884 and within four 

years annual shipments from the range exceeded one million 

tons. Detailed examination of the district began shortly afterward by the U.S. Geological Survey and culminated with a 

report by Irving and Van Hise (1892) that presented a remarkably complete and accurate model of the nature, control, and 

origin of the natural ores. Although refined by later works 

such as Hotchkiss (1919), the model was not substantially 

modified for more than a century and withstood perhaps the 

ultimate test—it guided further exploration that led to the 

discovery of hundreds of millions of tons of ore.

Natural ores occur exclusively in the Ironwood IronFormation as irregular masses of soft, earthy iron oxides and 

hydroxides. Iron-carbonate and iron-silicate minerals within 

the Ironwood oxidized and chert was replaced with iron 

hydroxides and oxides. In some ores, relict bedding was preserved. The ore bodies are mainly localized along the keels of 

plunging structural troughs, which are most commonly formed 

by the intersection of diabase dikes with the top of the Palms 

Formation or with the top of shaley units within the Ironwood. 

Figure 14 is simplified from Hotchkiss (1919, fig. 26) and 

summarizes the essential features of the geometry of the ore 

bodies. Iron ore extends from the surface to depths in excess 

of 1.5 km; the greatest depth of mineralization has not been 

determined. 

The origin of ore as a product of a structurally focused 

flow of deeply circulating, oxygenated ground water that 

both oxidized the original iron minerals and replaced the 

chert with iron minerals was first proposed by Irving and Van 

Hise (1892) and is still the accepted theory of the ore genesis 

mechanism. More than 70 years of mining and study of the 

Gogebic iron range has not produced any convincing evidence 

to contradict their original hypothesis. 

The age of mineralization is quite unconstrained. Irving 

and Van Hise (1892) correctly surmised that ores formed 

entirely after the rocks were in their present structural orientation so that dike-bed intersections formed plunging troughs. 

The structural tilting of the range occurred at about 1,060 Ma 

(Cannon and others, 1993), thus placing a lower limit on the 

age of mineralization. A similar constraint is suggested by the 

distribution of ore relative to the effects of Mesoproterozoic 

metamorphism. A relationship between metamorphic recrystallization and ore formation was noted by James (1955). 

Natural ores in all of the iron ranges in the Lake Superior 

region are restricted to areas of very low-grade metamorphism 

where the original very fine grain size was preserved. Even 

modest metamorphic recrystallization inhibited ore formation, apparently because the increase in grain size caused by 

metamorphic recrystallization diminished the grain surface 

area-to-mass ratio of individual iron minerals and made 

oxidation and dissolution less efficient. In the Marquette iron 

range, for instance, natural ores are confined to the eastern part 

of the range, which is in the chlorite zone of regional metamorphism, but they are absent elsewhere in the range, where, 

although geologically similar, the rocks are metamorphosed 

to biotite and higher grades (James, 1955). A similar relationship is present in the Gogebic iron range. The western limit of 

natural ores near Upson, Wis., coincides with the appearance 

Neoarchean

granite

Palms Formation (argillite)

Palms Formation (quartzite)

Natural ore

Diabase dike

Natural ore

Figure 14. Schematic block diagram showing 

the relation of natural ores to the stratigraphy and 

structure of the central Gogebic iron range. Simplified 

from Hotchkiss (1919). Ore bodies were localized in 

the downward-closing structures in the Ironwood 

Iron-Formation formed by intersection of diabase 

dikes and impermeable shale beds beneath or within 

the iron-formation.
Economic Geology 33

of metamorphic recrystallization in the contact metamorphic 

aureole of the Mellen Intrusive Complex. Natural ores are 

absent in the remaining western part of the range. The age of 

the Mellen Intrusive Complex is about 1,100 Ma (Cannon and 

others, 1996), which indicates that ore formation must have 

been younger, but there is virtually no upper limit on the age 

of mineralization.

Some aspects of ore formation are still poorly understood. There has been little modern interest in studying these 

ores, because the ore bodies have been largely depleted and 

are otherwise of little commercial interest. In addition, access 

to ore bodies is no longer available. Some interesting remaining questions are as follows:

1. Under what topographic, tectonic, and climatic conditions 

can vigorously circulating ground-water systems reach 

depths of more than one kilometer?

2. What was the chemistry of ground water that allowed 

large amounts of iron to be dissolved, transported, and 

redeposited under apparently strongly oxidizing conditions in which ferric iron is generally virtually insoluble?

Recently, Morey (1999) proposed that the high-grade iron 

ores of the Mesabi iron range in Minnesota, which are in some 

respects similar to ores on the Gogebic iron range, may have 

formed as a result of large-scale lateral movement of heated 

formation waters during Penokean deformation. According to 

his theory, the Mesabi ores were formed in Paleoproterozoic 

time and their origin is not related to the present erosion surface. The Gogebic iron range ores clearly seem to have been 

controlled by the Mesoproterozoic structures and metamorphism, and are therefore unlikely to have formed during Paleoproterozoic fluid flow as suggested for the Mesabi iron range.

Iron ores in all of the mining districts in northern Michigan and Wisconsin contain a suite of accessory minerals 

including adularia, barite, calcite, manganite, psilomelane, 

and rhodochrosite. These minerals typically occur in cavities 

within the hematite and goethite or limonite ores. The accessory minerals locally cement fragments of the iron ore. Their 

nature and occurrence suggest that they formed by the same 

processes that produced the iron ore. These accessory minerals have the potential to provide information as to the age and 

temperature of formation of the ores; however, no systematic 

study of these minerals has been undertaken. 

Taconite

Taconite has not been mined on the Gogebic iron range, 

but a large taconite resource has been identified in the western 

part of the range. The resource was quantified by Marsden 

(1978) in a study for the U.S. Bureau of Mines. Much of Marsden’s work was based on information gathered by U.S. Steel 

Corporation during the late 1950s (see previous discussion on 

the structure of the western Gogebic iron range), and on the 

projection of that information into less studied areas. Marsden 

(1978) estimated that a 35-km-long belt of taconite extended 

from Upson to Mineral Lake in Wisconsin. He divided the 

Ironwood Iron-Formation into five units that are somewhat 

different from the classic members established farther to the 

east. Based on metallurgical tests that indicated the percentage 

of recoverable iron as magnetite, he identified two continuous ore zones separated by a layer of waste rock. The lower 

ore zone included most of the Plymouth Member and the 

upper ore zone included mostly the Pence and Norrie Members. Marsden then did an economic simulation of the amount 

of material that could be mined and concentrated at a profit 

using economic assumptions from Mesabi iron range taconite 

operations. The principal economic constraint was the depth to 

which open-pit mining could proceed in these steeply dipping 

units before the cost of stripping the hanging-wall waste rock 

made deeper mining uneconomic. Marsden calculated that 

3,711,000,000 tons of taconite could be mined profitably with 

prices, costs, and technologies of the mid-1970s. This resource 

constitutes one of the largest undeveloped iron resources of the 

Lake Superior region. Economically recoverable amounts of 

ore may be substantially different from Marsden’s estimate of 

25 years ago and could be either larger or smaller depending 

on the interplay of value and cost of production. The previous 

estimates did not consider environmental and aesthetic factors 

of large open-pit mines and related processing facilities, which 

would undoubtedly be major factors affecting the future of this 

resource.

Parts of the Ironwood Iron-Formation in the eastern 

Gogebic iron range are also considered to be a taconite 

resource, but are much less surely quantified. The rather broad 

outcrop belts of mostly moderately dipping magnetic iron-formation from sec. 17, T. 47 N., R. 44 W., to sec. 22, T. 47 N., 

R. 43 W., seem to offer some possibility of future economic 

mining. Metallurgical data on recoverable iron is not available for these deposits so the percentage of the iron-formation 

that might contain acceptable grades of iron recoverable by 

magnetic separation cannot be estimated.
34 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

Synopsis

The Gogebic iron range is an integral part of the Paleoproterozoic continental margin assemblage, which was 

deposited on the south-facing margin of a craton composed 

of Archean crystalline rocks. The iron range formed near the 

inner edge of the foreland fold-and-thrust belt of the Penokean 

orogeny in which the rocks of the Marquette Range Supergroup were deposited and soon after were deformed. The 

rocks of the Gogebic iron range provide the most complete 

and best-studied record of events in this part of the orogen.

Sedimentation of the Marquette Range 

Supergroup

Environmental Setting of Deposition

Chocolay Group

The earliest recorded deposits of the Marquette Range 

Supergroup are clean quartz sands of the Sunday Quartzite, 

which is the basal unit of the Chocolay Group. Sedimentary 

structures indicate deposition in shallow water, at least partly 

in a tidal environment. The Sunday Quartzite, known only in 

the eastern part of the Gogebic iron range, becomes thinner 

toward the west and pinches out near the town of Wakefield, 

Mich. The extent to which this pinch-out is erosional rather 

than depositional is not known. The overlying Palms Formation lies unconformably on the Sunday so it is likely that the 

Sunday originally extended somewhat farther west than its 

current outcrop limit. The Sunday Quartzite grades upward 

into the Bad River Dolomite of the Chocolay Group, which 

indicates continuous deposition during which the clastic quartz 

supply was lost. The thick, partly stromatolitic, carbonate 

succession of the Bad River accumulated in shallow water. 

The Bad River is as much as 300 m thick in both the eastern 

and western parts of the iron range. In the eastern part of the 

iron range the Bad River always overlies the Sunday Quartzite, but in the west it directly overlies Neoarchean basement 

rocks and has only a few meters of arkosic dolomite at its 

base. The clean quartz sand of the Sunday Quartzite appears 

to be completely absent in the western part of the iron range. 

In the central part of the range, the Bad River is present only 

in a few isolated patches; however, the extensive chert breccia 

unit of the basal Palms Formation, which is interpreted to be 

a residuum of chert produced by dissolution of the Bad River, 

indicates that the Bad River Dolomite was probably continuous along most or all of the Gogebic iron range but was 

removed by erosion before the Palms Formation was deposited. In summary, the quartz sands of the Sunday Quartzite 

probably had a depositional pinch-out somewhere west of 

Wakefield, Mich. The Bad River Dolomite overstepped the 

pinch-out and was deposited as a continuous carbonate bank 

over the entire Gogebic iron range.

The time span between deposition of the Bad River 

Dolomite and the unconformably overlying Palms Formation 

is not well constrained, but may be substantial. Many authors 

have proposed that the Chocolay Group correlates with the 

upper part of the Huronian Supergroup of Ontario (Puffett, 

1969; Young, 1970; Ojakangas, 1982, 1985). If that is correct, 

then the Chocolay Group must be older than the 2,200 Ma 

Nipissing Diabase (Corfu and Andrews, 1986), which intruded 

rocks of the Huronian Supergroup. Recent age determinations 

for the Menominee Group and equivalent rocks (Fralik and 

others, 2002; Schneider and others, 2002) indicate and age of 

roughly 1,870 Ma. Thus, an erosion interval of more than 300 

my might separate the two groups.

Menominee Group

The deposition of the Menominee Group, which is the 

principal iron-bearing group of the region, began with finegrained, laminated muds of the lower Palms Formation. In the 

central part of the Gogebic iron range, the paleosurface was 

mantled with a residuum of chert fragments that formed by 

the dissolution of the Bad River Dolomite during the postChocolay Group erosion interval. This residuum was slightly 

reworked and mixed with quartz sand during the Palms transgression to form the basal chert breccia unit. Where the chert 

breccia is absent, the laminated argillite appears to lie directly 

on older rocks, which indicates that the base of the transgressive sequence formed in a relatively low-energy environment 

in which finely laminated muds were able to accumulate. The 

lower argillite grades upward into a medial unit of somewhat 

coarser grained, siliceous mudstone and sandstone, which 

in turn grades upward to thick-bedded quartzite (Ojakangas, 

1983). Although the total thickness of the Palms is somewhat 

variable along the length of the iron range, this pattern of 

coarsening-upward sediments is developed everywhere. The 

uniformity of the internal stratigraphy suggests that the Palms 

was deposited under stable conditions with no local influence 

of either a tectonically induced subsidence or uplift.

The upper quartzite of the Palms grades upward into the 

Ironwood Iron-Formation. The transition zone is between 1 

and 2 m thick and records the transition from the deposition 

of clean quartz sand of the Palms Formation to the chemical 

sedimentation of the Ironwood Iron-Formation. The Ironwood 

records the earliest basin instability in both the eastern and 

western parts of the Gogebic iron range. The central part of 

the iron range remained as a stable and quiescent depositional 

environment throughout Ironwood deposition. Water depth (or 

wave-base depth) varied enough to produce frequent alternations of quiet-water and agitated-water bottom conditions, but 

there is no evidence of either volcanism or syndepositional 

faulting anywhere in the central Gogebic iron range. In contrast, intrusive and extrusive igneous rocks are common in the 

eastern and western part of the iron range. Volcanic rocks are 

interbedded with the Ironwood and gabbro sills are common. 

The Presque Isle trough (discussed above) in the eastern part 

of the iron range, is a graben that formed during Ironwood 
Synopsis 35

deposition and controlled the thickness and character of the 

Ironwood and the interlayered Emperor Volcanic Complex 

(Klasner and others, 1998). The iron-rich depositional setting in which most of Ironwood accumulated eventually was 

transformed into a basin in which pyritic black shales were 

deposited. The black shales are preserved in some areas of 

the central part of the range and in the westernmost part of 

the range; however, they may have been more widespread 

but were removed by erosion before deposition of the Baraga 

Group.

The entire Menominee Group appears to be a transgressive sequence in which a barrier bar and back-bar lagoon 

transgressed over the low-relief surface developed on Neoarchean rocks and strata of the Chocolay Group (Ojakangas, 

1983). The fine-grained, laminated sediments of the lower part 

of the Palms Formation, which commonly overlie older rocks 

without a basal conglomerate phase, indicate that the lowenergy environment of a lagoon was the initial depositional 

setting of the transgression. Further transgression superposed 

the coarser quartzose sediments of the barrier bar over the 

lagoonal accumulations as shown by the coarsening-upward 

character of the Palms. The iron-rich facies that constitutes 

the Ironwood Iron-Formation apparently existed immediately 

seaward of the barrier bar and spanned the depth of wave base; 

therefore, the Ironwood was deposited directly on the coarse 

quartzose sand of the barrier bar. A euxinic basin, in which 

the pyritic black shale accumulated, appears to have been still 

farther seaward. With the continued transgression, the black 

shale was superposed over the iron-rich part of the Ironwood, 

which resulted in pyritic carbonaceous rocks in the upper part 

of the Ironwood and lower part of the Tyler Formation.

Baraga Group

The Baraga Group, consisting of the Tyler Formation in 

the west and the stratigraphically equivalent Copps Formation 

in the east, overlies the Menominee Group with a low-angle 

unconformable contact. The base of the Copps Formation is a 

conglomerate composed mostly of locally derived fragments 

of Neoarchean rocks and iron-formation. Likewise, the Pabst 

Member of the Tyler Formation is composed mostly of fragments of the underlying Ironwood Iron-Formation. The time 

span of the unconformity is not known, but it was sufficiently 

long enough to have allowed the Ironwood to become lithified 

before it was eroded and also long enough to have allowed 

erosion to completely remove the Chocolay and Menominee 

Groups in the eastern part of the Gogebic iron range. Roughly 

100 m of carbonaceous pyritic shale and cherty iron-carbonate-bearing iron-formation overlies the basal conglomerates. 

Most of the Copps and the Tyler is composed of graywacke 

and related clastic rocks that were deposited as a submarine 

fan. Detailed sedimentologic studies by Alwin (1976) established the turbidity flow character of this 3,000-m-thick section and documented a southerly source for the clastic material. Much of the detritus appears to have been derived from 

Archean crystalline rocks, which indicates that a significant 

upland of Archean rocks existed south of the Gogebic iron 

range at this time. At present, Archean crust extends only from 

the study area to the Niagara fault, about 50 km south of the 

Gogebic iron range, beyond which the crust consists entirely 

of Paleoproterozoic volcanic-arc rocks. Penokean tectonism 

may have greatly foreshortened this part of the continental 

margin.

Tectonic Setting of Deposition

The tectonic setting of deposition of the Marquette Range 

Supergroup has been discussed extensively (for instance, 

Larue and Sloss, 1980; Larue, 1981; Larue, 1983; Cambray, 

1978; Hoffman, 1987). The interpretation that deposition was 

in a foredeep basin, as proposed by Hoffman (1987), is consistent with much of the observational evidence (particularly for 

rocks of the Menominee and Baraga Groups) derived from the 

Gogebic iron range and other iron ranges of the region.

Chocolay Group

The interpretation that deposition of the orthoquartzites 

and dolomites of the Chocolay Group was on a stable platform 

or passive margin also has been widely accepted. Questions 

remain regarding the relation of the platform and passivemargin sequence to the subsequent depositional, volcanic, and 

deformational events. There are no radiometric constraints on 

the age of deposition of the Chocolay Group other than the 

age of the Neoarchean basement and the 1,870 Ma age for 

the unconformably overlying Menominee Group. A definitive solution to the age problem is not yet at hand, but the 

commonly proposed correlation of the Chocolay Group with 

parts of the Huronian Supergroup in Ontario is likely to be 

correct. This correlation requires the Chocolay Group to have 

been deposited before 2,200 Ma, which is a minimum age 

for the Huronian Supergroup (Corfu and Andrews, 1986). If 

the Chocolay Group is, in fact, older than 2,200 Ma, and if it 

formed during the continental breakup leading to the eventual 

Penokean collision at 1,870 Ma, then the passive-margin phase 

of the Penokean orogenic cycle must have persisted for more 

than 300 my.

Menominee Group

Radiometric ages for the Menominee Group (Schneider 

and others, 2002) indicate that the Menominee was deposited 

contemporaneously with the development and accretion of 

volcanic arcs in the Wisconsin magmatic terranes to the south. 

That relationship strongly suggests a foredeep depositional 

setting for the Menominee. The depositional model of Hoffman (1987) proposed that iron-formations were deposited 

on the outer ramps of foredeep basins in areas where normal 

faulting (created by crustal flexure) is common. With time, 

migration of the foredeep basin toward the continent superposed turbidite deposits of the axial zone of the foredeep 

basin over the iron-bearing sequences that were deposited on 
36 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

the outer ramp. Observations from the Gogebic iron range 

indicate that the Palms Formation was deposited in water that 

was shallow enough to record tidal fluctuations throughout its 

stratigraphic extent (Ojakangas, 1983). The overlying Ironwood Iron-Formation (at least in the central part of the range) 

also was deposited in water that was shallow enough to be 

periodically above the wave base. There are no indications that 

water depths were substantially below the wave base in any of 

the Menominee Group rocks in the central part of the Gogebic 

iron range. On the eastern and western ends of the range, the 

Palms Formation maintains its characteristic internal stratigraphy and evidence of a tidal depositional setting; however, 

the Ironwood Iron-Formation was deposited (at least in part) 

in subsiding fault-bounded troughs and was interlayered with 

volcanic rocks.

The Menominee Group in the central part of the Gogebic 

iron range appears to have been deposited in the most distal 

parts of the foredeep basin on the slowly and rather uniformly 

subsiding crust. Water depths were always shallow and fluctuated from tidal to somewhat below the wave base. Rocks 

on the eastern and western ends of the range appear to have 

been deposited in deeper water in the proximal parts of the 

foredeep basin, where more intense crustal flexing resulted in 

syndepositional normal faulting. A modified foredeep basin 

model was proposed by Schneider and others (2002) in which 

oblique collision of arc terranes accentuated crustal extension 

in the continental margin. This additional extension, when 

superposed on the extensional features that might be expected 

solely from orthogonal collision, might account for the welldeveloped depositional grabens and the abundant volcanic 

and mafic intrusive rocks immediately adjacent to tectonically 

stable parts of the basin.

The uniformly shallow-water depositional setting for the 

central part of the Gogebic iron range suggests deposition in 

the most continentward part of the foredeep basin, possibly 

north of a flexural forebulge (outer arch of Hoffman’s (1987) 

model). If so, the northward migration of the forebulge may 

account for the unconformity between the Menominee and 

Baraga Groups. 

Baraga Group

Within its lowest 100 m, the Baraga Group passes from 

a basal conglomerate, through euxinic black shale and cherty 

carbonate iron-formation, into deposits of a turbidite fan. The 

Baraga Group records the deposition of outer ramp facies 

(consisting of shales and iron-formation) followed by the 

deposition of a thick turbidite succession (Tyler and Copps 

Formations) in the axial zone of the foredeep basin. The Tyler 

is composed mostly of granitic detritus from a highland of 

Archean basement rocks that existed at that time to the south 

of the Gogebic iron range, and was deposited by northwardflowing turbidity currents (Alwin, 1976). Tectonic foreshortening of the continental margin is suggested by the roughly 

50 km distance between the Gogebic iron range and the margin of accreted arc terranes to the south. The foreshortening 

may have resulted in telescoping of the outermost extensions 

of continental crust during collision, which then elevated the 

area and allowed for eventual erosional destruction of thrust 

sheets that consisted of Archean basement rocks.

Volcanism

The chemistry of volcanic rocks also provides information on the tectonic environment in which the Marquette 

Range Supergroup was deposited. The Emperor Volcanic 

Complex is one of several volcanic units interbedded in the 

Marquette Range Supergroup. Other Paleoproterozoic volcanic 

units include the Clarksburg Volcanics Member of the Michigamme Formation, and the Hemlock Formation, and Badwater Greenstone. Detailed geochemistry, however, is available 

only for the Hemlock Formation and Badwater Greenstone 

(Ueng and others, 1988; Schulz, unpub. data [1984]). The 

composition of basalts from the Emperor Volcanic Complex 

compares very well with the composition of rocks from the 

Hemlock Formation and Badwater Greenstone. Both the 

Hemlock and the Badwater are predominantly subalkaline 

tholeiitic basalts that are characterized by moderately enriched 

light REE and other incompatible elements, and by variably 

negative Nb and Ta and positive Th anomalies (fig. 15). Unlike 

the Emperor Volcanic Complex and Badwater Greenstone, 

the Hemlock Formation is bimodal in that it locally contains 

abundant rhyolite of largely crustal origin (Schneider and 

others, 2002). The presence of crustally derived rhyolites in 

the Hemlock Formation is further evidence for the interaction 

between basaltic magmas and Archean basement during the 

evolution of the Menominee Group.

Determining the tectonic affinity of volcanic suites based 

on geochemical tectonic discrimination diagrams must be 

done with caution and in conjunction with knowledge of the 

geology of the area. This is particularly true when considering possible continental basalts as they often do not plot as 

expected in the within-plate basalt fields on commonly used 

tectonic discrimination diagrams (Holm, 1982; Wang and 

Glover, 1992). This results from the fact that some continental 

basalts show variable levels of relative depletion in high-fieldstrength elements (HFSE), particularly Nb, Ta and Ti, which 

form the basis of discrimination in many tectonic discrimination diagrams (Wang and Glover, 1992). The relative depletion 

of HFSE in some continental basalts, while still a subject of 

controversy, is typically attributed (1) to the derivation of the 

basalts from HFSE-depleted subcrustal lithosphere, and (or) 

(2) result from crustal contamination (Norry and Fitton, 1983; 

Dupuy and Dostal, 1984; Carlson, 1991).

The composition of the Emperor basalts, as well as that 

of the Hemlock and Badwater, resembles continental tholeiitic 

basalts; all three plot mainly in within-plate basalt fields on 

tectonic discrimination diagrams (fig. 16; see also Ueng and 

others, 1988). On some tectonic discrimination diagrams, 

the Emperor and related Paleoproterozoic basalts extend into 

arc-related basalt fields (for example, fig. 16D) because of the 

variable degrees of crustal contamination, which results in the 
Synopsis 37

relative depletion in Nb and Ta and an enrichment in incompatible elements such as Th and La .

The interpretation of the Emperor and related Paleoproterozoic basalts as continental tholeiites is compatible with the 

geology of the Menominee Group. The sequence lacks arcrelated rocks, which typically include (1) calc-alkaline andesites, and (2) large volumes of reworked volcaniclastic material 

that are typically part of volcanic suites in island-arc terranes. 

The Emperor and related basalts are either interbedded with 

or are laterally equivalent to relatively thick accumulations 

of fine-grained epiclastic and chemical sedimentary rocks, 

which is an association that suggests deposition in extensional 

basins far from an active magmatic arc. Schulz and others 

(1993) proposed that the Menominee Group and its contained 

volcanic rocks were deposited during the active rifting phase 

of the evolution of the Paleoproterozoic continental margin in 

the Lake Superior region. A more recent study of rhyolite from 

the Hemlock Formation yielded a U-Pb zircon age of 1,874±9 

Ma (Schneider and others, 2002). This date suggests that (1) 

the Hemlock Formation and correlative volcanic units of the 

Menominee Group were coeval with arc-related volcanic and 

plutonic rocks that are now preserved in the accreted terranes 

to the south in Wisconsin (Sims and others, 1989), and (2) that 

the Menominee Group is in a second-order basin in a foredeep 

related to oblique subduction and subsequent collision along 

the Superior craton margin (Schneider and others, 2002).

A possible analogue of the proposed Menominee Group 

setting is the Late Silurian to Early Devonian rift that was 

caused by dextral transpression in the northern Appalachians 

(Dostal and others, 1993; Malo and others, 1995). Volcanic 

rocks of mostly Middle Silurian to Early Devonian age locally 

constitute a significant part of the supracrustal sequences that 

unconformably overlie Ordovician and older rocks that were 

deformed during the Ordovician Taconian orogeny (Keppie, 

1992). The volcanic sequences are bimodal in that they contain (1) continental tholeiite to transitional basalts and 

(2) rhyolites that were derived by crustal anatexis (Dostal and 

others, 1989, 1993; Van Wagoner and others, 2002). Basalts 

from the Late Silurian to Early Devonian rift basins in the 

northern Appalachians have a composition that is similar to 

that of the Paleoproterozoic basalts of the Emperor Volcanic 

Complex; both types exhibit variably negative Nb and Ta 

anomalies (fig. 17). 

The sedimentologic, structural, and geochemical 

evidence favors the interpretation that the Menominee and 

Baraga Groups were deposited along a convergent margin that 

lay north of the accreting arc terranes that are now prevalent 

south of the Niagara fault. Subsidence and uplift were largely 

flexural responses of the southern edge of the Archean craton 

that was being overridden by volcanic arcs. Extensional features such as the Presque Isle trough, which were important in 

localizing volcanic rocks and controlling thickness and facies 

variations in sedimentary rocks, were second-order extensional 

features caused by the oblique collision of the southern edge 

of the craton with the volcanic arcs of the Wisconsin magmatic 

terranes.

100

10

1

ROCK/PRIMITIVE MANTLE

Th Nb Ta La Ce Nd Zr Hf Sm Ti Gd Tb Dy Y Ho Er Tm Yb Lu V Sc

Basalts of the Hemlock Formation

and Badwater Greenstone

EXPLANATION

Emperor Volcanic Complex

Diabase near Atkins Lake, 

 Wisconsin

Figure 15. Primitive-mantlenormalized multi-element plot 

showing the compositional 

similarity of mafic rocks from 

the Emperor Volcanic Complex 

and diabase near Atkins Lake 

(symbols same as in figure 7) with 

basalts of the Hemlock Formation 

and Badwater Greenstone 

(shaded field). Hemlock data from 

Ueng and others (1988); Badwater 

data from Schulz (unpublished 

data). Primitive-mantle 

normalizing values from Kerrich 

and Wyman (1997).
38 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

Figure 16. Tectonic discrimination diagrams showing the within-plate affinity of the mafic rocks of the Emperor Volcanic Complex and 

diabase from near Atkins Lake. (A) from Pearce and other (1977). (B) from Pearce and Norry (1979), (C) from Pearce and Cann (1973), and 

(D) from Cabanixs and Thiéblemont (1988). 

A B

C D

EXPLANATION

Emperor Volcanic Complex

Diabase near Atkins Lake, 

 Wisconsin

Continental

Ocean island

FeO(t)

Al MgO 2O3

Within-plate

Island arc

Ti/100

Zr Y*3

Back-arc

basin

Anorogenic

Tb*3

Th Ta*2

Island arc

Mid-ocean

ridge basalt

(MORB)

Within-plate

1

10

100

10 100 1,000

Zr (ppm)

Zr/Y

Ocean ridge

Orogenic

Ocean floor

Clac-alkaline

D A

B

C Calc-alkaline

Arc tholeiite

N-MORB

E-MORB

Ocean island-continental alkaline

Continental

Orogenic tholeiite

Penokean Deformation

The intensity of Penokean folding and faulting varied 

substantially from the central part of the Gogebic iron range 

to both the eastern and western parts. In the central part of 

the range, folds are absent except for small folds near faults. 

Faults formed in at least three sets. Early, bedding-parallel 

faults were well documented in mine workings, but are no longer accessible for study. The best developed of these faults was 

traced for the entire length of the productive part of the iron 

range. The amount of displacement was not well determined 

but was probably small, in the range of hundreds of meters. 

These faults were most logically north-directed thrust faults, 

but kinematic indicators from the fault zones have not been 

described. Many more such faults may occur in other parts of 

the central part of the range but have not been detected. Slatey 

cleavage that formed locally in the argillaceous beds of the 

Tyler Formation (especially in higher parts of the section) suggests that a significant northward transport may have occurred. 

Two sets of steep faults have offset the bedding-parallel faults. 

Their contribution to Penokean deformation is not completely 

resolved; at least some of the faults were reactivated during 

Mesoproterozoic time or later.

In the eastern part of the range, the effects of Penokean 

deformation appear abruptly at the Sunday Lake fault near 

Wakefield, Mich. This reverse fault resulted in about 2 km of 

offset of the bedrock units and was truncated by the unconformity at the base of the Mesoproterozoic strata, indicating 

that it is clearly a Penokean structure. East of the Sunday Lake 

fault, Penokean folds and faults are widespread. Folds plunge 

to the east or west when restored to their Paleoproterozoic 

orientation. Faults were of sufficient magnitude to produce 
Synopsis 39

structural repetitions within the iron-bearing sequence. The 

relations here show that there was a sharp transition from 

very weakly deformed rocks of the central part of the range to 

much more intensely folded and faulted rocks of the eastern 

part of the range.

In the western part of the range, a similar sharp transition 

occurs between the nonfolded and only moderately faulted 

rocks as far west as the Tyler Forks River and the tightly 

folded and thrust-faulted rocks farther west. The structural 

pattern, as shown in map view, is deceptively simple west 

of the Tyler Forks River. Mesoproterozoic block rotation of 

the Penokean fold-and-thrust belt resulted in a map view that 

approximates a cross section drawn parallel to fold axes. This 

geometry inherently does not indicate the true geometry of the 

folds. When the map pattern is visualized without the Mesoproterozoic rotation, it reveals tight folds and a significant 

repetition of rocks by thrust faults. As in the east, restoring 

structures to their Paleoproterozoic orientations indicates that 

fold axes plunge gently to the east or west. Thrust faults are 

100

10

1

ROCK/PRIMITIVE MANTLE

Th Nb Ta La Ce Nd Zr Hf Sm Ti Gd Tb Dy Y Ho Er Tm Yb Lu V Sc

Silurian to Devonian basalts

Basalts of the

Emperor Volcanic

Complex

Figure 17. Primitive-mantlenormalized multi-element plot 

showing the compositional 

similarity of rift-related Late 

Silurian to Early Devonian 

basalts from the northern 

Appalachians (symbols) and 

the mafic rocks from the 

Emperor Volcanic Complex 

and diabase near Atkins Lake 

(shaded field). Data for Late 

Silurian to Early Devonian 

basalts from Dostal and 

others (1989, 1993). Primitivemantle normalizing values 

from Kerrich and Wyman 

(1997).

largely parallel to bedding, although numerous ramps also 

were mapped.

The Paleoproterozoic rocks of the Gogebic iron range 

illustrate significant facies changes across a boundary that represents a change from relatively stable deformational conditions (the central part of the range) to more tectonically active 

conditions (eastern and western parts of the range). This same 

boundary was also a deformational front during Penokean contractional deformation, which suggests that Penokean deformation was controlled in part by older structures that formed 

during crustal extension or earlier.

Finally, Paleoproterozoic and Neoarchean rocks were 

tilted northward as rigid blocks within the upper plate of a 

crustal-scale listric thrust fault at about 1,100 Ma (Cannon and 

others, 1993). Thus, the entire Gogebic iron range, in present map view, represents a longitudinal cross section of the 

Penokean fold belt, which provides a unique view of Paleoproterozoic sedimentation and deformation within the Penokean 

orogen.
40 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

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Acknowledgments

The authors wish to acknowledge the efforts of several 

individuals whose assistance was vital to the completion of 

this report. Cheryl Hedgeman assisted with much of the field 

work in mapping the western part of the Gogebic iron range 

from 1991 through 1994. Connie Dicken and Stacy Saari 

performed the digital cartographic compilation for plate 1 

and many other illustrations. The original manuscript was 

improved by thorough constructive technical reviews by 

Avery A. Drake, Jr., and Suzanne W. Nicholson.

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44 The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt

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Manuscript approved for publication July 17, 2006.

Prepared by the Reston and Columbus Publishing Service Centers.

Edited by Elizabeth D. Koozmin.

Graphics and cartography by A.D. Long.

For more information concerning the research in this report, please 

contact William F. Cannon, U.S. Geological Survey, MS 954, National 

Center, 12201 Sunrise Valley Dr., Reston, VA 20192.

Cannon and others— The Gogebic Iron Range—A Sample of the Northern Margin of the Penokean Fold and Thrust Belt—Professional Paper 1730

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