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Produced by the NASA Center for Aerospace Information (CASI)
s

MR

a akR14 4 A;

FINAL REPORT

i May 1979 to 31 May 1980

on

Agreement NCC,5-13

between

NASA

and

The Department of Earth and Planetary Sciences

The Johns Hopkins University

Baltimore, Maryland 21218

"Petrologic and Geophysical. Study of the Source of

Long Wavelength Crustal Magnetic Anomalies"

yK ect Director

Bruce D. Marsh

Professor

Dept. of Earth & Planetary Sciences

The Johns Hopkins University

Baltimore, ?Maryland 21218

CorKeL^p ndertt

Charles M. Schlinger

Assistant Professor

Dept. of Geology & Geophysics

University of Utah

{ Salt Lake City, Utah 8012

(NASA -C.R-17.5168) PETROLOGIC AND GEOPHYSICAL 1484-20946.

STUDY OF THE SOURCE OF .LONG WAVELENGTH

IJ CRUSTAL MAGNETIC ANOMALIES Final. Report., 1

May 1979 - 31 May 1980 (Johns Hopkins wi.v. ), Unclas

'19 p HC A02/MF A01 CSCL 08N G3/46 12520
t.

Introduction

' t

This is the final report for NASA grant NCC 5-13 entitled:

Petrologic and Geophysical Study of the Source of Long Wavelength Crustal

Magnetic Anomalies, beginning 1 May 1979 for one year. The summary of

what was originally proposed is given below from the abstract of the

original, proposal.

Original Abstract

In this report we document our progress on "the petrologic and geophysical study of the source of long wavelength crustal magnetic anomalies":

NASA grant NCC5-13. We also request continuation of funding for a three year

i

period. We have made substantial progress in the study of these anomalies

by analysis of both aezomagnei:ic and satellite data; large remanent F

magnetizations have been identified as sources of several anomalies. We

have made a quantitative investigation of POGO satellite orbital errors and

their manifestation as pseudo-anomalies, and have synthesized a large body of

I literature in the geparate disciplines of rock magnetism, iron-titanium oxide

E geochemistry, continental structure, composition, and evolution, and the

petrology of igneous and metamorphic rocks (with an emphasis on their Fe-Ti

oxide minerals). This syntheses, we feel, is essential in any attempt of

data interpretation and source description.

For the new award period we propose theoretical, experimental, and field

study of the formation Fe-Ti oxides, their occurrence in crustal rocks, and

their consequent_ magnetic propezties during progressive metamorphism. We plan

i
i

to study the petrography and magnetic properties of samples collected

in the vicinity of the large Mauritanian anomaly. A diversified study

of this breadth has not been attempted, do date. Finally, we propose to

determine the effect(s) of oxygen fugacity (fq2) on both crystallization

and p4rtial melting (e.g. granulite facies metarmorphism) of crustal

rocks and their constituent Fe-Ti oxides. As a result, we hope to uncover the process which allows, through normal geologic processes and

conditions, Fe-Ti oxides to produce these anomalies.

This report is in two parts: (I) The magnetic mineralogy and

magnetic signature of banded iron formations; and (II) Phase equilibria of

magnetite in orogenic magmas.

Part I

The Magnetic Mineralogy and Magnetic Signature of Bana4.. iron Formations

In an analysis of either satellitemagnetic or aeromagnetic data one

may interpret anomalies, to first approximation, in terms of ideally shaped

highly magnetic sources distributed in the crust. From anomaly intensity

and signature it is possible to deduce the appro:K,amate source magnetization 

and geometry. The use of further geologic and geopl%ysical constraints 

serves to produce a more credible model. Iron formations of Precambrian

age are known to exr;end laterally for hundreds of kilometers, however,

their average thickness is no more than one kilometer (usually lass than

several hundred meters) and their average magnetization is poorly knor,,*n.

Here we consider the magnetic mineralogy of iron formations and their

n . ' _, -. ... - s ^ •

ice + d,^/

l^i.y..^^ t ^.. ..' ..-3. —

- - .wr^.r..^-^1:
}

signature in aeromagnetic data.

The worldwide occurrence of vast banded iron formations (BIF) of

sedimentary origin in the Precambrian. rncord remains an enigma. Once

it was thot.tght that such deposits were coeval with ages clustered at

about 1.9 to 2.5 B.Y. (Goldich, 1973). It has recently been proposed

(Cole & Klein, 1981) that BIFs are in fact distributed throughout much

of the Precambrian, that they are not coeval. The observation remains

that the world's largest BIFs are indeed Proterozoic with similar age

dates as proposed by Goldich.

Large reserves of iron ore in BIFs are known on every continent.

Of the world's greatest deposits those of Lake Superior, U.S.; the

Labrador trough, Canada; Krivoy Rog & Kursk, U.S.S.R.; Minas Gerais,

Brazil; and the Hammersley Basin, W. Australia are perhaps best known.

These deposits are all approximately contemporaneous in age (James

& Sims, 1973; Dorr, 1973). The modern thickness and extent of these

deposits are•often unrelated to their primary geometry as they've often

suffered alteration by tectonic deformation by folding and thrusting,

erosion, and granitization (Dori, 1973). The largest BIF deposits known,

such as those of the Hammersley group and Frere formation of W. Australia

it the Caue itabirite of the Quadrilatero Ferrifexo, Minas Gerais,

Brazil were laid down as extensive blanket deposits. In Brazil the

Caue itabirite is thought to have once covered 100,000 k.2 with an average

thickness of 250 m (Dori, 1973). The Hammersley group may have once

extended over 150,000 km2 with average thickness of 1000 m (Trendall, 1973).

Trendall & Blockley (1970) estimate an iron content in the Hammersley basin

r :.s .,." --^ 1.^ ^. ,^a^.-i"...^i• 1I•-,ice ..r-, ^Y ^` x sK^'. m,:.._^.. _...__.^^^
of 10 20 g. The Labrador trough deposits extend almost continuously for

over 800 km (Dimroth et.a?., 1970) and the Kuruman & Penge BIF of South

Africa extend discontinuously for 950 km (Beukes, 1973). The width of the

South African and Canadian BIF is an order of magnitude less than their

linear extent.

Banded i g on formations are distinguished on behalf of the mkneralogy,

or "facies", of their constituent iron-bearing minerals, according to

James (1954):

1) Oxide facies, e.g., magnetite, hematite, or both

2) Sulfide facies, e.g., pyrite

3) Carbonate facies, e.g., siderite, ankerite (FeMg(CO3)2)

4) Silicate facies, e.g., Greenalite (Fe-chlorite), Minnesotaite

(Fe-talc), Stilpnomelane (layered K-Fe-phyllo silicate),

Fayalite, Grunerite (MgFe7 Si8022 (OR) 2 ) or orthopyroxene.

The facies classification is critical for interpretation of the depositional

environment of BIF (e.g., James, 1954; Eugster & Chou, 1973; Kimoereley,

1978; Gross, 1973, 1980).

We are more interested in the mineralogy of these formations,

principally in the oxide and silicate facies, as it relates to their

magnetization, hence, the influence of metamorphism is of great importance

as most BIFs have been metamorphosed to some degree, The magnitization of

BIFs exposed today at the earth's surface reflects their respective diageneti

and metamorphic history; this is'a'iso ;;rue for unknown formations deep in

the crust. Thus, observations of BIFs of differing metamorphic grade will

allow us to interpret their mineralogy in terms of consequent magnetization.
To properly examine this issue we will consider the metamorphic mineralogy

of deposits from all geologic ages. 

Diagenetic (unmetamorphosed) and Low Grade Banded Iron Formations

In the literature there is no consensus on the primary mineralogy

of BIM's; most have been metamorphosed to some degree and I will avoid a

detailed discussion of this controversy. In a study of very low grade

Archean BIF from W. Australia Cole (1980) reported samples from all facies

of James • (1954). His relevant overvations include:

1) Absence of hematite in fresh (unweathered) sample,

2) Ubiquitous appearance of magnetite, in highly variable amounts

but secondary in abundance to Minnesotaite ('Fe3Si4ko10(OH)2),

siderite and quartz.

3) Magnetite is never observed in ,reaction textures.

French-(1973_ p. 1069) claims that "virtually all of the magnetite in

the Biwabik iron-formations and other iron-formations in the L. Superior

region is relatively coarse grained and secondary (La Berge, 1964; French,

1968; Han, 1972) and forms by replacement of pre-existing silicates and

carbonates". The oxidation of siderite to magnetite (Klein, 1973) by the

reaction

6FeCO3 + 02 -} 2Fe 304 + 6CO2

is thought to explain the petrographic evidence for such replacement

(La Berge, 1964). In a compilation of diagenetic and low-grade metamorphic
4

rl

replacement textures, (French, 1973) magnetite is seen to be present in

secondary assemblages but only on orie occasion is it replaced by a

carbonate, ankerite. Klein & Fink (1976) in their study of very low-grade

assemblages within the Proterozoic Sokoman iron-formation (Howells river

P.rea, Labrador trough) found:

1) Magnetite occurs in primary form, recrystallized after a fine

grained magnetite or hydromagnetite (Fe 'i) 4 .nH2O) precursor. This

interpret. . d origin is contested by Floran & Papike (1978) and

French (1973) --see below.

2) Hematite occurs in primary form.

3) Magnetite bearing facies are the most common in the sedimentary

column: both silicate and hematite bearing facies also contain

magnetite. 'Magnetite is pervasive in the silicate, oxide,

carbonate facies (Klein & Bricker, 1977)°

4) The sulfide facies is of minimal Extent in the column.

In a study of the 2 b.y. old contact metamorphosed Gunflint ironformation, Minnesota, Floran & , Papike (1978) conclude:

1) Magnetite is ubiquitous in all zones havi-,, formed by reduction

of hematite.

2) Where magnetite is involved in prograde metamorphic reactions

such as the disappearance of greenalite, Fe 6Si40 10 (OH) 8 , by reactions

Greenalite + 1/202 -} Minnesotaite + Fe304 + 3H20

Greenalite + 1/20 2 -> 2Fe304 + 4SiO2 + 4H20

it is always a product and never a reactant..
Alternatively Klein (1978) documents the reaction

Greenalite + 4Si0 2 + Minnesotaite + 2H20

in the regionally metamorphosed Labrador trough.

High-Grade Mineralogy

Gole & Klein (1981a) described the high-glade mineralogy of Archean

BIFs in W. Australia. The occurrence of coexisting clinopyroxene, orthopyroxene, and fayilite allow estimation of pressure and temperature

usin 2-pyroxene and olivine-orthopyroxene geothermometers (T = 670±50 0C,P = 3-5k

They found:

1) "Magnetite is not a major constituent... at any of the locations".

Even so they state that magnetite "is abundant in quartz rich .

assemblages and is a minor constituent of most other... assemblages".

Unfortunately, no modal measurements are presented. In a review of

South American iron-formations Dom (1973) notes that iron rich bands in

itabirite contain hematite at lower metamorphic grades and magnetite at

higher metamorphic grades. Sulfide facies iron-formations while volumetrica11y minor contains pyrite throughout low to moderate-grades of

metamorphism, however, under high-grade conditions pyrrhotite is instead 

present (Gunderson & Schwartz, 1962; French, 1968). 

. The mineralogy of a 13ingle iron-formation.exhibiting assemblages

ranging from essentially unmetamorphosed variants up to transitional

amphibolite-granulite facies equivalents has been reviewed by Klein (1978).
Iron-formations of the Labrae,.)r trough extend for nearly 700 miles in a

north-south direction; in the south unmetamorphosed iron-formation grades

into the Grenville front where high-grade assemblages are to be found.

The range of P-T conditions varies from 150 0C at 1-2 kb in the north to

700-7500C at 10-11 kb in the south; relevant observations, interpretations,

and conclusions include (Klein, 1978):

1) "The hihest grade iron-formation assemblages contain abundant

orthopyroxene, magnetite, carbonates, some almardine and only

minor amphiboles. This is a common assemblage in several other

iron-formations of high-grade origin".

2) Quartz, hematite, and magnetite are ubiquitous in samples ranging

from very low-grade to high-grade, however, samples may contain

only hematite or only magnetite. Magnetite and hematite occasionally

coexist and their textures imply equilibrium along the Fe304-Fe203

buffer.

The behavior of magnetite and hematite in banded iron-formations

throughout metamorphism can now be summarized. In the oxide facies bol,:h

minerals are stable over the entire range of metamorphic pressures and

temperatures; locally oxidizing or reducing conditions will determine their

relative stability or coexistence. Clearly, if hematite is initially

present, the reduction to magnetite will require the presence of one or

more of free carbon, CO, or H 2 (Eugster, 1972). In the silicate, oxide, 

and carbonate facies oxidation or decarbonation of siderite during

diagenesis may yield magnetite (IIan, 1972; Beukes, 1973; Klein, 1:973;

French, 1968).

+"" ^ .. ;,^,. :y . ,./Yi-3.rte`. i^ i- - ''^ 2 c"w: in ^.T.u..^+.^cfa
I r

In general magnetite appears to be if not primary then a .secondary or

replacement product of diaaenetic and metamor hhic reactions. Additionally

the sulfide facies undergoes transition from non-magnetic pyrite bearing

assemblages to magnetic pyrrhotitic acrt.rimblages in response to high-groie

metamorphism. From field studies of many of the world's greatest ironformations we have evidence indicating the stability of magnetite over

hematite in highly metamorphosed locations (Darr, 1973). Deposits with

hematite are thought to result from near-surface or surface oxidation of

magnetite (Tolbert, et al., 1973; Dorr, 1973; DeVilliers, 1971). Such

near-surface oxidation and ore enrichment processes generally produce

hematite and martite (hematite pseudomorph of magnetite and, sometimes,

pyrite) at the expense of magnetite. At any metamorphic grade there

is a considerable quantity of Fe bound by silicate and carbonate phases;

we cannot assume that the majority is contained within the oxides. In

spite of this observation the magnetic mineralogy of unaltered to highly

metamorphosed iron-formation at the surface or as yet undiscovered at depth

can be interpreted as consisting of magnetite, hematite, and on occasion,

pyrrhotite. Measurement of relevant magnetic properties would be useful

for our general knowledge yet given the blanket or through nature of BIF,

however laterally extensive, we may tentatively conclude that only the

largest deposits might be crustal sources of regional magnetic anomalies.

Also, it seems to be the case that metamorphism does enhance the magnetization

of banded iron-formations whatever their type facies, sulfide, carbonate,

silicate, or oxide.

A crude estimate of the potential magnetization of the larger banded

iron-formation provinces such as the Hammersley basin of W. Australia can
be made using Trendall & Blockley's (1970) estimate of total Fe content

in the Hammersley basin: ti10 22 grams. Typical estimates of source

dimensions of larger regional magnetic anomalies such as that over the 

Central. African Republic are ti1000 km by N100 km by N20 km thick. Modeling

(e.g., Regan & Marsh, 1982) indicate, a magnetic susceptibility for such

a source region of ti10

..2

 c.g.:s. (12.56X10 2 SI). Such a susceptibility can 

be attained with 3% by volume of magnetite, or ti6X1018cm3 , or equivalently, 

N3X10 19g. Multiplying this mass of magnetite by the ratio, molecular

weight of Fe/molecular weight of Fe30

V

 one finds the equivalent mass of

Fe for such mass of magnetite. This turns out to be v2.3X10 19g of elemental

Fe, Assuming Trendall & Blockley (1970) estimate to be correct in order

of magnitude we can speculate that if only one quarter to one half of this. 

mass of Fe is bound by magnetite, then such an iron formation could serve

as a source of a regional magnetic anomaly. A similar estimate for the

iron formations of the Kursk magnetic anomaly is given by Alexandrov (1973),

of 5X10 19g. In the Quadrilatero Ferrifero of Minas Gerais, Brazil, an

intense regional, magnetic anomaly is found (Bosum, 1973); the interpretation

is that this feature is caused by the immense iron formation of that

district. It is very difficult to prove any of these interpretations as 

we are working with a tremendous slab of crust, QX10 20cm3 . Even so we

may speculate that if one of these vast iron formations has been metamorphosed to medium or high-grade under moderately reducing conditions

(relative to the Fe30i} Fe203 buffer) and structurally reconfigured into

a true and dimensional body it could serve as the source of a more intense

regional magnetic anomaly such as that of Kursk, the Quadrilatero Ferrifero,

or perhaps even that of Central Africa.
1

Part 11

Phase Equilibria of magnetite in Orogenic Magmas

The principal carrier of magnetization in crystal rocks is magnetite.

The abundance of this phase in plutonic rock depends critically on the order

of its appearance in the sequence of crystallization. If magnetite appears

early in the sequence it may be modally significant, while if it appears

late in the sequence most of the iron may have been taken up by pyroxenes

and olivine and magnetite will be sparse. The position of magnetite in

this sequence depends critically on the fugacity (fO 2 ) of oxygen in the

system. A high f0 2 promotes magnetite stability while a low f02 reduces

its thermal stability relative to the silicate minerals. This stability

has been previously shown by R. N. Thompson for an iron-rich rock, and

his results are shown by Figure 1.

This composition is not particularly similar to that of orogenic

magmatic bodies which largely make up the crust. And we originally proposed

to carry out a similar study on a basaltic composition from an active island

arc; the Aleutian island arc. The method we used consisted of sealing the

powdered rock in an evacuated silica glass tube containing a chemical to

buffer the f02 . The charge was then heated to magmatic temperatures in a

horizontal resistence furnace and held at that temperature for several days.

Counting the runs necessary in developing the technique, thirty four runs

were made over the period of one year.
t

aj l

The starting composition in all runs is:

AT-6

Si02

49.05

TiO2

0.89

Al20320.32

Fe2 03

3 -.,,

FeO 6.52

MnO 0.18

M80 3.91

CaO 10.44

Na20 3.36

0.89

^N205

.^_.27 0

Total 99.14

The buffering materials were Ni-NiO, Co-CoO, and MnO-Mn 304 . Each of these

pairs fixes the fO 2 to vary with temperature along a curve that is similar

to what is actually observed in magmas. The charges were quenched from

magmatic temperatures, the rock-glass powdered and inspected under a microscope

in oils to determine the presence or absence of magmetite. These results

are shown by Figure 2 each symbol represents two runs, solid = magnetite,

open = no magnetite.

It is clear from these results that magnetite stability is indeed

strongly affected by fO 2 . The critical range of fO2 is in the region slightly

above the Ni.-NiO buffer. To determine the stability limit of magnet ite in
in an actual magma, the prevailing magmatic fO2 must be known.

The determination of fO2 in a magmatic assemblage is most commonly

done using chemical thermodynamics and less commonly through direct

sampling with an xygen probe. Basalts from the ocean ridges typically

have an fO2 near that described by the quartz-magnetite-fayal.it .e (QF'M)

reaction. Continental lavas like rhyolites have an fO 2 near that described

by the MnO-M304 reaction, and island arc lavas have art f02 in between these

two extremes. I+la fact, orogenic (island arc) lavas generally have phenocrysts of magnetite, while ocean ridge lavas rarely do.

The importance of this effect cannot be overemphasized in terms of

its importance to magnetization of the crust. A magma crystallizing at

depth under these conditions will produce an abundance of magnetite.

This magnetite will be coarse-grained and will aquire a weak remanent

magnetization upon cooling through the Curie point. The magnetic susceptibility of the resulting rock is directly proportional to its magnetite

content as is shown by Figure 3.

Conclusions

The stability of magnetite in crustal rocks is intimately dependent on

the sequence of crystallization of the original magma as well as subsoli.dus

metamorphic reactions. The position of magnetite in the sequence of

crystallization depends to a great extent on the prevailing fugacity of

oxygen, which is fairly high. Because of this, phenocrysts of magnetite

can be expected early in the crystallization sequence. Should these rocks

form a large part of the lower continental crust they could easily amount

for much of its magnetization.
ORIGINAL PACE 19

OF POOR QUAL17Y

1

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'^Y~~ ^^°^' ^b& ~ W.i~_~.SMs. no`
a • ,

s 

. 16 

ORIGINAL PAGE 69

OF POOR QUALITY

f02-Temperature plot for one-atmosphere melting experiments

of a tholeiite 

-8.5

-9..0

log fog

-9.5

-10.0

V

Mt•!L

F1+Mt+L /ENO

O1+P1+

Mt+L /.e^ Liquid

Pl

/ +

/ 01+ L

P1+L /

140 1060 1080 1A00 1120

Temperature (OC)

Figure 1 ` The f02-Temperature plot for one-atmosphere melting experiments on a 

tholeiite (after Thompson, 1973). 'L is liquid, P1 is plagioclase,

Mt is magnetite, 01 is olivine. NNO is the Nickel-Nickel Oxide buffer

curve and QFM is the Quartz-Fayalite-Magnetite buffer curve (after

Ewart, 1979)
s ,

r ^

17

omta1igAL PAGE 10

OF POOR QUALI'T'Y

&ckj A r- 6

z 

3

4

^ 7

^• 8

I

q

to

e^

'0 ^4-rp'; ""-

^t^t

a^

VrP -

o _

.. l 

0

0 1I00 l^40 ^ o

O C

Figure 2: Experimental results obtained under

this grant on the stability of

magnetite in the orogen:c composition

AT-6.
-e c-,

0 c

O r- E

E '"E r= a o o.,

a =u! r- .

—---------

to

I I Jit I I LL-L-.L-LLLL

O CIL

cr

01X 'AgjiqijdzmftS

0 ciW

1:1

.w 0

H z

0 0

u p

Co w

iJ

-W 41

to u

co

44 'r4 0

C3

n I a

o

^^

14 -W 0 -H

> 3

In C $4

O

O

0 ;

♦ ^l ..............

O

v:

E ;a

O

o a •

^r o

cz

M

C! I

c;

O

ORIGINAL PAC;r. 

OF POOR QUALI

O