Geochronology and geochemistry of Grenvillian igneous suites in the northern Oaxacan Complex, southern Mexico: tectonic implications

Geochronology and geochemistry of Grenvillian igneous suitesin the northern Oaxacan Complex, southern Mexico: tectonic

implications

J. Duncan Keppie c,*, J. Dostal a, K.L. Cameron b, L.A. Solari c,F. Ortega-Gutierrez c, R. Lopez d

a Department of Geology, St. Mary’s University, Halifax, NS, Canada B3H 3C3b Department of Earth Sciences, University of California, Santa Cruz, CA 95064, USA

c Instituto de Geologıa, Universidad Nacional Autonoma de Mexico, Ciudad Universitaria Mexico, Delegacion Coyoacan, 04510 Mexico

DF, Mexicod Geology Department, West Valley College, Saratoga, CA 95070, USA

Accepted 18 October 2002

Abstract

Chemical and U�/Pb isotopic analyses of metaigneous rocks in the northern Oaxacan Complex in southern Mexico

indicate that they form part of two granitic�/gabbroic suites intruded at �/1157�/1130 and �/1012 Ma, which were

metamorphosed under granulite facies conditions between �/1004 and 980 Ma. Although the older suite has both

within-plate and arc geochemical signatures, the arc characteristics (enrichment of La and Ce relative to Nb, Ta, and

Th) are inferred to result from crustal contamination, a conclusion consistent with their negative oNd signatures. The

younger suite is spatially associated with anorthosites (from which we were unable to acquire a protolith age),

suggesting that collectively it forms part of anorthosite�/mangerite�/charnockite�/granite (AMCG) suites. The tholeiitic

nature of the mafic rocks along with the within-plate character of the felsic rocks suggests that they were intruded

during extension related to either farfield backarc rifting, rifting above a slab window, or anorogenic intercontinental

rifting. Potentially correlative AMCG suites are widespread in Mexico, the Grenville Province of eastern Canada and

northeastern USA, and the Andean massifs of Colombia, however, Pb isotopic data most closely resemble those in

South America. These data are consistent with published hypotheses that suggest Oaxaquia represents an exotic terrane

derived from Amazonia.

# 2002 Elsevier Science B.V. All rights reserved.

Keywords: Geochronology; Geochemistry; Igneous; Grenvillian; Oaxaquia; Mexico

1. Introduction

The �/1 Ga rocks of the Oaxacan Complex of

southern Mexico represent the largest exposure

(10 000 km2) of the basement of the Oaxaquia* Corresponding author. Fax: �/52-5-622-4303

E-mail address: [email protected] (J.D. Keppie).

Precambrian Research 120 (2003) 365�/389

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0301-9268/02/$ - see front matter # 2002 Elsevier Science B.V. All rights reserved.

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composite terrane (Fig. 1). Oaxaquia is defined as

a Precambrian�/Paleozoic composite terrane that

extends beneath Mesozoic and Cenozoic rocks

from the Ouachita Orogen along the backbone of

Mexico to southern Mexico (Ortega-Gutierrez et

al., 1995), and may extend into the Chortis block

of Honduras, which may have lain south of

Mexico until the Eocene (Schaaf et al., 1995;

Keppie and Ortega-Gutierrez, 1999). Thus, the

combined size of Oaxaquia and the Chortis block,

1 000 000 km2, makes it important in reconstruc-

tions of Rodinia.

There are generally two schools of thought

about the location of Oaxaquia/Chortis in Rodinia

reconstructions (Fig. 2). One school believes that

Oaxaquia represents a southern continuation of

the Grenville Orogen of eastern and southern

Laurentia (De Cserna, 1971; Shurbert and Cebull,

1987) forming a connecting segment between

Laurentia and either Antarctica (Moores, 1991;

Dalziel, 1992), the Albany�/Fraser�/Musgrave

belts of Australia (Brookfield, 1993; Karlstrom et

al., 1999; Burrett and Berry, 2000), or Siberia

(Sears and Price, 2000; Fig. 2, location #1). The

other school infers that Oaxaquia is an exotic

terrane that originally lay off either eastern

Laurentia or Amazonian, and was transferred to

southern Laurentia in the Permo-Carboniferous

during the formation of Pangea (Fig. 2, location

#2 or #3; Keppie, 1977; Ballard et al., 1989; Yanez

et al., 1991; Keppie and Ortega-Gutierrez, 1995;

Ortega-Gutierrez et al., 1999; Keppie and Ortega-

Gutierrez, 1999; Keppie and Ramos, 1999; Ramos

and Aleman, 2000; Cawood et al., 2001; Keppie et

al., in press). On the other hand, Ruiz et al. (1999)

have proposed a split model based upon whole-

rock Pb isotope data, in which they divide

Oaxaquia along the Trans-Mexican volcanic belt

and correlate the northern part with Texas, and

the southern part with Colombia. On the other

hand, Cameron et al. (2002) show that feldspar Pb

isotope data for Oaxaquia fall on a linear array

(distinct from Texas) that probably result from

mixing of two end-members, which supports the

concept of Oaxaquia as a single block. This paper

presents the results of geochronological and geo-

Fig. 1. Location of the Oaxacan Complex of southern Mexico in relation to Oaxaquia.

J.D. Keppie et al. / Precambrian Research 120 (2003) 365�/389366

chemical analyses of igneous rocks from the

northern part of the Oaxacan Complex (Fig. 3),

which bear on this problem.

2. Geological setting

Prior to our recent studies, research indicated

that the northern Oaxacan Complex consists of

paragneisses (marbles, calcsilicates, quartzofelds-

pathic and graphitic gneiss) and orthogneisses

(anorthosite, charnockite, amphibolite, and peg-

matite; Anderson and Silver, 1971; Ortega-Gutier-

rez, 1984), that were involved in a single folding

event (Kesler and Heath, 1970; Kesler, 1973)

accompanied by granulite facies metamorphism

with peak temperatures and pressures of 700�/

825 8C at 7.2�/8.2 kb under restricted PH2O

conditions (Mora et al., 1986). The oldest dated

charnockite yielded a U�/Pb age of �/1113 Ma

(Silver et al., 1994) with pegmatites giving con-

cordant U�/Pb zircon ages of 10509/20, 9759/10,

and 9609/15 Ma (Anderson and Silver, 1971;

Ortega-Gutierrez et al., 1977). K�/Ar cooling

Fig. 2. One Ga reconstruction of Rodinia showing the three possible locations (1�/3) for Oaxaquia (OX) and the Chortis block (CH)

modified after Keppie and Ramos (1999), Dalziel et al. (2000) Sears and Price (2000). Inset is an enlargement of location #3 (modified

after Dalziel, 1994); �/1 Ga orogens are shaded and stars indicate the locations of AMCG complexes. Abbreviations: A, Arequipa

massif; CA, Carolina and Goochland terranes; CM, Coats Land/Maudheim/Grunehogna terrane; D, Dalradian of Scotland and

Ireland; E, Ellsworth�/Whitmore Mountains; M, Moine of Scotland; R, Rockall Plateau; SF, Sao Francisco craton; SM, Santa Marta

massif; SO, Sunsas Orogen.

J.D. Keppie et al. / Precambrian Research 120 (2003) 365�/389 367

ages on hornblende, muscovite, biotite and K�/

feldspar from cross-cutting pegmatites yielded

�/927, 925, 875, and 775 Ma, respectively (Fries

et al., 1962; Fries and Rincon Orta, 1965).

Fig. 3. Geological map and structural section of the northern Oaxacan Complex.


Our recent studies (Solari et al., 2002) havedocumented that the northern Oaxacan Complex

may be divided into a series of thrust slices (from

bottom to top; Fig. 3): (#1) anorthosite, gabbro-

norite, jotunite, and monzodiorite; (#2) cumulitic

mafic gneiss and charnockite; (#3) migmatitic

orthogneiss; (#4) paragneiss intruded by metasye-

nite, charnockite, and anorthosite. Anorthositic

dykes intrude the migmatitic orthogneiss suggest-ing limited relative displacement between units 1

and 3 (Solari et al., 2002). On the other hand,

larger displacement across the Phanerozoic, E-

vergent, phyllonitic thrust zone between units #3

and #4 is suggested by the lack of clear correla-

tion. These studies indicate that the northern

Oaxaca Complex underwent a complex tecto-

nothermal history involving migmatization at �/

1100 Ma (Olmecan event) and granulite facies

metamorphism and deformation between �/1004

and 980 Ma (Zapotecan event; Solari et al., 2002).

All of the slices and structures were refolded by

upright-steeply inclined, NNW-trending folds dur-

ing the Phanerozoic. Early Ordovician rocks that

rest unconformably upon the northern Oaxacan

Complex were deformed by N-trending uprightfolds of latest Paleozoic�/early Mesozoic age

(Centeno-Garcia and Keppie, 1999). U�/Pb geo-

chronological analyses presented here indicate at

least two episodes of intrusion for most plutonic

igneous units: ]/1130�/1157 Ma (generally upper

intercept ages for mafic rocks, metasyenite and

charnockite) in the upper thrust slice and �/1012

Ma (anorthosite, mafic cumulatic gneiss, gabbro-norite, jotunite, garnet-bearing charnockite) in the

lower two thrust slices.

3. Petrography

Several phases of regional metamorphism of

Precambrian age have been preserved in the rocks

of the northern Oaxacan Complex. An oldermigmatization event is restricted to the thrust slice

#3, but was mineralogically replaced by younger

regional granulite facies metamorphism that is

pervasive throughout the complex. On the other

hand, amphibolite to greenschist facies recrystalli-

zation preferentially affected the lower slices and

localized shear zones. The high-grade metamorph-ism and deformation is responsible for the general

granulitic textures and banded or foliated appear-

ance of the rocks.

3.1. Upper thrust slice (#4)

In the abundant quartzofeldspathic lithologies

of the upper slice this metamorphism produced the

entire orthopyroxene series of granitic rocks(charnockites to enderbites), which are character-

ized by the strongly perthitic (mesoperthite) and

antiperthitic nature of the feldspar, the common

presence of garnet and the rutilated nature of

much of the quartz, which gives a bluish to

purplish aspect to many of these rocks. Hornble-

nde is very common, especially in the syenitic units

where the amphibole coexists with quartz andmesoperthite with or without ortho- and clinopyr-

oxene. The typical, massive, green charnockite

contains alkali feldspar�/quartz�/orthopyroxene�/

clinopyroxene�/hornblende�/ore9/garnet�/biotite.

Mafic lithologies in the upper slice are char-

acterized by granoblastic to strongly foliated,

hydrous and anhydrous parageneses with ortho-

pyroxene, clinopyroxene, plagioclase, hornblende,and opaque ore. The following assemblages were

found in textural equilibrium: plagioclase�/

orthopyroxene�/hornblende�/phlogopite9/ore, pla-

gioclase�/clinopyroxene�/hornblende�/ore, clin-

opyroxene�/plagioclase�/ore, orthopyroxene�/

hornblende�/plagioclase�/phlogopite�/ore, plagi-

oclase�/garnet�/orthopyroxene�/hornblende, plagi-

oclase�/orthopyroxene�/clinopyroxene�/hornble-nde, and very rarely the high pressure assemblage

plagioclase�/garnet�/clinopyroxene�/quartz�/

orthopyroxene�/ore. Garnet and quartz are both

present only in rocks where orthopyroxene and/or

ilmenite reacted with plagioclase to form the high

pressure assemblage clinopyroxene�/garnet�/

quartz. Biotite and phlogopite generally form

late mineral coronas around ilmenite or occupyirregular spaces between the other phases. The

abundant and widespread hornblende and mica

were probably stable at granulite facies due to

their high content of titanium or by a strongly

magnesian composition in intermediate rocks.

Garnet without quartz not uncommonly coexists


with the two pyroxenes and hornblende, and couldhave been stabilized by a more aluminous or

ferrian compositions. Plagioclase has extinction

angles and a negative sign indicating a composi-

tion in the andesine range in most of the mafic

rocks. An anorthosite body in the upper thrust

slice at Union Zaragoza consists mainly of plagi-

oclase with minor orthopyroxene, perthite, clino-

pyroxene and quartz and accessory apatite, ore,biotite and garnet.

3.2. Lower thrust slices (#1 and #2)

Pure anorthosite in the lowest thrust slice has a

granoblastic texture composed of antiperthitic

andesine with apatite, ilmenite, and rarely quartz

as accessories. Antiperthite may be coarsely lamel-lar with up to 30% orthoclase exsolution. Asso-

ciated mafic and felsic gneisses in the lower two

thrust slices are characterized by anhydrous as-

semblages and garnet upto several centimeters in

size and forming upto 60% of the rock. Most mafic

minerals were altered during retrogression, which

obscures nature of the granulitic assemblages.

Nevertheless, two-pyroxene garnetiferous assem-blages predominate and a single sample yielded

inverted pigeonite, probably inherited from the

original magmatic mineralogy. As in the upper

slices, the high pressure assemblage clin-

opyroxene�/garnet�/quartz was very rarely devel-

oped.

Retrogressive metamorphism formed spectacu-

lar simple, double and triple coronas of amphi-boles, micas, and epidote around pyroxenes and

titaniferous hornblende. Anthophyllite, cumming-

tonite and tremolite are the most common amphi-

boles together with uralite in the inner coronas,

whereas brown or green biotite and epidote

formed the outer rims. The retrogressive meta-

morphism proceeded into the greenschist facies

with chlorite, epidote, calcite and white micaforming veins and patches throughout these rocks.

Talc preferentially replaced orthopyroxene as

pseudomorphs. Most alkali feldspar was converted

to microcline, whereas white mica, epidote, calcite

and leucoxene are alteration products in the

anorthosite.

4. Geochronology

U�/Pb geochronological analytical methods fol-

low procedures outlined in Lopez et al. (2001).

Seven samples were selected for isotopic analysis:

gabbro, metasyenite, and charnockite from the

upper thrust slices, and anorthosite and associated

meta-gabbronorite, mafic gneiss and charnockite

from the lower thrust slices (Fig. 3). All sampleshave been affected by Grenvillian granulite meta-

morphism and associated deformation at �/980�/

1004 Ma (Solari et al., 2002).

4.1. Upper thrust slice (#4)

Zircons from the metagabbro (sample #66A98)

are generally colorless and clear, and range in

shape from elongate with aspect ratios of 1:2 (Fig.4b) to equant to tabular. Cathodoluminescence

reveals that many of the grains have complex

internal zoning that may be igneous in origin (Fig.

4a: image 66A98); however, all zircons are multi-

faceted (Fig. 4b) and this presumably reflects

metamorphic modification and/or growth. The

five analyzed fractions are 1�/2% discordant but

have 207Pb/206Pb ages that range from 1145 to1077 Ma (Table 1). Two youngest fractions are

abraded whereas the three oldest are not; thus in

this sample at least, there is no evidence that

abrasion either decreases discordance or selectively

removes the metamorphic component of the

zircons. There is no systematic age distinction

between the tabular and elongate zircons (no

equant crystals were analyzed). A regression ofthe five fractions yields intercepts of 12579/71 and

10219/39 Ma (Fig. 5a) with a probability of fit of

0.21. The lower intercept lies within error of the

age of the Grenville event, and the poorly con-

strained upper intercept is interpreted as the

minimum age of crystallization of the gabbro.

Charnockite (sample 5998) was collected within

the uppermost paragneiss thrust slice (#4). Itcontains both equant and elongate (aspect ratios

as high as 1:4) zircons. Individuals from both

populations showed igneous growth zoning (Fig.

4a: image 5998), and no correlation was found

between morphology and age (Table 1). Six single

zircons were analyzed, and the results illustrate


some of the problems common in interpreting

geochronological of the Mexican Grenvillian gran-

ulites. Four of the six analyzed fractions are

concordant within analytical uncertainty; how-

ever, their Pb�/Pb ages differ by more than 100

million years (fractions 1a and 1b). The two oldest

Fig. 4. (a) Cathodoluminescence images of zircons from the northern Oaxacan Complex (white bar on each image corresponds to 100

mm scale); (b) photomicrographs for some of the analyzed zircons (scale is provided in each image).


Table 1

U�/Pb geochronological data for samples of the northern Oaxacan Complex, southern Mexico

Description% Weight

(ug)

U

(ppm)

Total

Pb

(ppm)

Com.

Pb (pg)

206Pb/204Pb

raw dataa

Atomic ratiob Age (Ma)b %Dis.

207Pb*/206Pb*

%error

207Pb*/206Pb*

%error

208Pb*/206Pb*

%error

206Pb*/238U

%error

206Pb*/238U

%error

207Pb*/235U

%error

207Pb*/235U

%error

206Pb*/238U 207Pb*/235U 207Pb*/206Pb*

66A98, Gabbro Host

1, dm, elong.

(1)

19 327 63 12 3381 0.07792 0.13 0.08226 0.19199 0.19 2.0627 0.24 1132 1137 11459/3 1.1

2, dm, eqm (4) 21 205 39.7 46 1108 0.07729 0.09 0.06368 0.18892 0.20 2.0133 0.22 1116 1120 11299/2 1.2

3, dm, flat, (6) 13 177 33.3 19 1479 0.07761 0.14 0.05573 0.18875 0.35 2.0199 0.38 1115 1122 11379/3 2.0

4, dm, 1:2, abr,

(4)

15 151 28.4 37 702 0.07601 0.20 0.04766 0.18265 0.44 1.9143 0.48 1081 1086 10959/4 1.3

5, dm, flat, (2) 23 242 44.2 28 2226 0.07532 0.07 0.08351 0.17989 0.19 1.8682 0.20 1066 1070 10779/2 1.0

5998, Charnockite

6, dm, elong,

abr, (1)

22 143 29 4 8848 0.07840 0.21 0.13203 0.19606 0.23 2.1195 0.31 1154 1155 11579/4 0.3

7, dm, equant,

abr, (1)

9 182 39 33 617 0.07839 0.22 0.14594 0.19565 0.25 2.1147 0.34 1152 1154 11579/4 0.4

8, dm, equant,

abr, (1)

8 218 44 19 1161 0.07793 0.15 0.10073 0.19476 0.45 2.0927 0.48 1147 1146 11459/2 �/0.2

9, dm, equant,

abr, (1)

15 184 33 6 5445 0.07508 0.07 0.08884 0.17889 0.16 1.8518 0.18 1061 1064 10719/2 0.9

10, dm, elong,

abr, (1)

17 254 45 8 5981 0.07412 0.07 0.08944 0.17603 0.11 1.7990 0.13 1045 1045 10459/2 0.0

11, dm, elong,

abr, (1)

16 251 45 21 4186 0.07350 0.15 0.14556 0.17004 0.28 1.7233 0.32 1012 1017 10289/3 1.6

6098, Syenitic Granulite

12, dm, elong,

(2)

64 92 19 33 2034 0.07636 0.07 0.19232 0.18170 0.17 1.9132 0.18 1076 1086 11059/2 2.6

13, dm, 1:2,

abr, (6)

87 65 13.6 14 4824 0.07714 0.06 0.20047 0.18928 0.16 2.0133 0.17 1117 1120 11259/2 0.7

14, dm, 1:2,

abr, (1)

37 68 15 35 863 0.07713 0.41 0.213127 0.18931 0.44 2.0133 0.61 1118 1120 11259/8 0.6

15, dm, flat,

(2)

44 82 17.3 53 788 0.07813 0.14 0.171668 0.18523 0.33 1.9954 0.36 1095 1114 11509/3 4.8

16, dm, elong,

gb, (1)

16 86 19 36 455 0.07873 0.24 0.21166 0.18544 0.60 2.0129 0.64 1097 1120 11659/5 5.9

6498, Mafic Granulite

17, elong, abr

(7)

26 77 15 39 409 0.07270 0.33 0.12624 0.16501 0.21 1.6540 0.39 985 991 10059/7 2.1

18, elong, abr

(3)

29 54 9 28 2938 0.07255 0.17 0.11245 0.16368 0.38 1.6373 0.42 977 985 10019/4 2.4

19, elong (1) 61 44 8 33 1671 0.07265 0.20 0.11348 0.16668 0.19 1.6696 0.28 994 997 10049/4 1.0

20, elong, abr

(1)

14 93 16 21 1281 0.07280 0.25 0.09469 0.16747 0.74 1.6812 0.77 998 1001 10099/5 1.0

OC9810, Charnockite

21, elong, abr

(1)

42 105 17 24 8771 0.07263 0.12 0.04608 0.16854 0.18 1.6878 0.21 1004 1004 10049/3 0.0

22, elong, abr,

gb (1)

19 90 17 38 482 0.07254 0.38 0.11871 0.16798 0.65 1.6802 0.77 1001 1001 10019/8 0.0

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Table 1 (Continued )

Description% Weight

(ug)

U

(ppm)

Total

Pb

(ppm)

Com.

Pb (pg)

206Pb/204Pb

raw dataa

Atomic ratiob Age (Ma)b %Dis.

207Pb*/206Pb*

%error

207Pb*/206Pb*

%error

208Pb*/206Pb*

%error

206Pb*/238U

%error

206Pb*/238U

%error

207Pb*/235U

%error

207Pb*/235U

%error

206Pb*/238U 207Pb*/235U 207Pb*/206Pb*

23, elong, (1) 15 100 17 10 1676 0.07285 0.12 0.08150 0.16758 0.51 1.6834 0.52 999 1002 10109/3 1.1

6398, meta-Gabbronorite

24, elong, 1:3,

abr, (3)

39 82 14 23 1426 0.07264 0.24 0.088666 0.16441 0.34 1.6466 0.41 981 988 10049/5 2.3

25, elong, 1:2,

abr, (1)

17 77 13.46 24 592 0.07292 0.30 0.069865 0.16702 0.69 1.6793 0.75 996 1001 10129/6 1.6

26, elong, abr,

(1)

43 65 11 20 1438 0.072143 0.19 0.07729 0.16590 0.33 1.6502 0.38 989 990 9909/4 0.1

6298, Anorthosite

27, frag, abr,

(10)

83 32 6 35 800 0.072574 0.19 0.19235 0.16407 0.50 1.6418 0.54 979 986 10029/4 2.3

28, frag, abr,

(8)

93 30 6 21 1375 0.072525 0.12 0.29647 0.16543 0.33 1.6542 0.35 987 991 10019/3 1.4

*, Denotes radiogenic Pb; %, dm, diamagnetic; abr, abraded; rnd, round; stby, stubby; elong, elongate; (1:2), aspect ratio; number in parenthesis is number of grains

analyzed. All fractions were analyzed using the 205Pb:235U mixed isotopic tracer.a The highly radiogenic 206Pb:204Pb ratios coupled with the small sample size of many fractions documents the low Pb background in our lab.b Decay constants used 238U�/1.55125�/10�10; 235U�/9.48485�/10�10, 238U/235U�/137.88. Errors on the U:Pb ratio used for plotting on the concordia diagram are

9/0.4% based on replicate analyses of Geostandard 91500 and other zircons (see Lopez et al., 2001). Errors on the U:Pb ratio are in percent (%), and calculated using the

program PBDAT (Ludwig, 1991) with two sigma errors on the measured isotopic data. The calculated errors shown are generally better than 0.4%, which indicate the

analytical precision of our lab. The 207Pb*:206Pb* age uncertainties are two sigma and also from the program PBDAT (Ludwig, 1991). Total processing Pb blank amount

varied between 2 and 30 pg during the course of this study, but were usually B/10 pg. Initial Pb compositions are from feldspar separates from the dated samples.

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Fig. 5. U�/Pb analyses of zircons plotted on concordia diagrams from the northern Oaxacan Complex: (a) metagabbro (sample

#66A98); (b) metasyenite (sample #6098); (c) charnockite (sample #5998); (d) garnet�/clinopyroxene mafic gneiss (sample #6498); (e)

charnockite (sample #OC9810); (f) meta-gabbronorite (sample #6398); (g) anorthosite (sample #6298).


fractions are concordant or essentially so, andhave 207Pb/206Pb ages of 11579/4 Ma (Fig. 1). This

is a minimum age for the sample, but we interpret

it as near the crystallization age because the207Pb/206Pb age was replicated. We interpret re-

maining fractions to have been disturbed by a

granulite metamorphic event at �/1004�/980 Ma.

A regression line with a forced lower intercept of

990 Ma passes through or at least touches theerror ellipses of all six fractions, including fraction

1b that is concordant at 10459/2 and fraction 2

concordant at 11459/2 (Fig. 5b). We cannot

completely rule out the possibility that the con-

cordance of these two points has geologic signifi-

cance; however, there is no independent evidence

to support that interpretation.

Metasyenite (sample 6098) was also collectedwithin the paragneisses of slice #4. It contains

clear, colorless, equant, tabular, and elongate

(aspect ratios as high as 1:2) zircons that show

igneous growth zoning (e.g. Fig. 4a: image 6098).

The analyzed zircons have rather complex U�/Pb

systematics with four fractions having similar207Pb*/235U ages between 1114 and 1120 Ma but

widely differing 207Pb/206Pb and 206Pb*/238U ages(Table 1). Fractions 9 and 10 (Fig. 4b) have

identical 207Pb/206Pb ages and are the least dis-

cordant (B/1%). A regression of those two and

fraction C yields an upper intercept of 11319/10

Ma that we interpret to be near the crystallization

age of the sample (Fig. 5c). Fractions 12 and B,

which have older 207Pb/206Pb ages, may have

contained an inherited component that experi-enced Pb loss.


Garnet �/orthopyroxene mafic gneiss (sample

#6498) from thrust slice #2 immediately overlying

the anorthosites contains extremely abundant

zircons with as many as 50 visible in a singlethin-section. Most are elongate with aspect ratios

as high as 1:4, and some of the grains appear to

have igneous zoning (Fig. 4a: images 6498a and b).

A regression of the four analyzed fractions yields

an upper intercept of 10129/12 Ma (Fig. 5d) with a

probability of fit of 0.78. Due to the apparent

igneous zoning, we interpret the upper intercept to

represent the time of intrusion.

Charnockite (sample OC9810) was collected

from a 100 m thick layer within the mafic gneiss

of thrust slice #2. The zircons are tabular to

elongate (aspect ratios as high as 1:3), and some

show zoning, which may be igneous but is more

ambiguous than in the samples discussed pre-

viously (Fig. 4a: image OC9810). Of the three

analyzed fractions, two are concordant with207Pb/206Pb ages of 10049/3 and 10019/8 Ma

(Fig. 5e, Table 1), and we interpret these as

metamorphic ages. Fraction D, which has a207Pb/206Pb age of 10109/3 Ma may contain a

component of protolith zircon.

Metagabbronorite (sample 6398) was closely

associated with anorthosite in the lowest thrust

slice. Most zircons from this sample are elongate

with aspects ratios of about 1:2�/1:3 (Fig. 4b). One

of the three fractions analyzed is concordant with

a 207Pb/206Pb age of 9909/4 Ma (Fig. 5f). The two

remaining fractions are about 2% discordant and

have somewhat older 207Pb/206Pb ages. It is temp-

ing to interpret the age of the fraction concordant

9909/4 Ma as that of granulite facies metamorph-

ism, but the morphologies of all the analyzed

fractions were similar (Fig. 4b).

Anorthosite (sample 6298) contains zircons that

are irregular in shape and rounded, and the

crystals lack clear igneous zoning (Fig. 4a: image

6298). Regression of the two analyzed points yields

an upper intercept of 9999/9 Ma (Fig. 5g), and we

believe this age is dominated by the granulite

metamorphic component.

These results indicate that the older igneous

units (�/1100 Ma) are located in the upper two

thrust slices. In contrast, there are no 207Pb/206Pb

ages or upper intercept ages from the lower two

thrust slices older than 1012 Ma, and it is difficult

to resolve the protolith igneous and metamorphic

age components of these rocks. Sample 6498

shows fairly clear igneous growth zoning (Fig.

4a) and its upper intercept age of 10129/12 Ma

maybe close to the protolith age. The zircons from

the remaining samples from the lower thrust slices

are probably dominated by the metamorphic

component.


5. Geochemistry

Samples were analyzed for major and trace

elements by X-ray fluorescence at the Regional

Geochemical Center at St. Mary’s University,

supplemented by those on the dated samples,

which were performed at the University of Cali-

fornia at Santa Cruz. The precision and accuracy

of the data have been reported by Dostal et al.(1986, 1994). Representative samples were then

chosen for the analysis of rare-earth elements

(REE), Th, U, Ta, Zr, Nb, and Y by inductively-

coupled plasma-mass spectrometer at Memorial

University of Newfoundland. The analytical error

of the trace element determinations is 2�/10% and

for the major elements is B/5%. Representative

whole-rock analyses are given in Table 2.

5.1. Upper thrust slices (#3 and #4)

Igneous rocks in the upper thrust slices crop out

in the northern part of the area and may be

subdivided into intermediate and mafic rocks

(charnockites and metagabbros).

5.1.1. Charnockites

These rocks (OX-9-11, OX-15, OX-63-67, C-42)

are mafic to intermediate in composition with SiO2

ranging from 48 to 63 wt.%, and straddle al-

kaline�/subalkaline boundaries on the Zr/TiO2

versus SiO2 (Fig. 6) and alkalies versus SiO2

diagrams. All the rocks are quartz-normative,

have low MgO, Ni, Cr (Table 2) and Mg#

(MgO/MgO�/FeOtot in mole%; 0.21�/0.45). Theyare, however, high in K2O (3�/5 wt.%) and plot

into the high K or shoshonitic field on a K2O

versus SiO2 diagram (Fig. 7). They also contain

elevated concentrations of Ba (�/2000 ppm), Sr

(300�/700 ppm), Zr (500�/900 ppm) and Nb (�/25

ppm) and have high Ti/V ratios (65�/85; Fig. 8).

Several major elements correlate with SiO2: TiO2,

P2O5, MgO, CaO and CaO/Al2O3 decrease whileK2O and total alkalies (Na2O and K2O) increase

with increasing of SiO2 (Fig. 7). K2O/Na2O ratio is

typically �/1 and correlates positively with SiO2

(Fig. 7a). The decrease of CaO, MgO and CaO/

Al2O3 with increasing differentiation (as exempli-

fied by SiO2 increase) is indicative of crystal-

lization dominated by clinopyroxene. Antithetic

variations of P and Ti with SiO2 imply the

fractionation of apatite and Fe�/Ti oxides, respec-

tively. The REE patterns of these rocks are

subparallel and enriched in light REE (LREE;

Fig. 9a), with Lan �/125�/300 and (La/Yb)n �/5�/

10. Those with high REE contents show a slight

negative Eu anomaly. The mantle normalized

trace element patterns of the rocks are also

subparallel and display a distinct enrichment in

La and Ce relative to Nb, Ta and Th, and in Ba

relative to Rb and Th (Fig. 10a). In general, the

rocks possess many similarities to recent shosho-

nites and also resemble the orthopyroxene series

(charnockite�/mangerite to quartz mangerite) of

the anorthositic suites (Rock et al., 1988; Wyman

and Kerrich, 1989; Conradie and Schoch, 1988).

Unfortunately, their relationship to the Union

Zaragoza anorthosite, whose age has not been

determined, is presently unknown. Compared with

shoshonitic lamprophyres and shoshonites, the

rocks have low Th contents and Mg# (Owen et

al., 1992). On the other hand, some of the felsic

rocks have some alkaline arc affinities, such as

enrichment of La and Ce relative to Nb, Ta, and

Th, and in Ba relative to Rb and Th. Although this

signature is characteristic of volcanic arcs, it may

also be the result of crustal contamination (Forster

et al., 1997). The latter is suggested by the negative

oNd values of the Oaxacan charnockites (ca. �/1.5;

Patchett and Ruiz, 1987; Ruiz et al., 1988). The

low Th content and Th/La ratio of the rocks

probably reflect the effect of granulite facies

metamorphism that generally leads to a depletion

of Th (Dostal and Capedri, 1978; Rudnick and

Taylor, 1987).

According the classification of Muller and

Groves (2000) based upon Zr/Al2O3 versus TiO2/

Al2O3 plot, the rocks were emplaced in a within-

plate setting. The samples also plot in the within-

plate granite field on Nb�/Y and Rb�/(Y�/Nb)

diagrams (Fig. 11). In this respect, they are similar

to many Grenvillian plutons in the Grenvillian

magmatic belt that have within-plate composi-

tional affinities (Owen et al., 1992). Like many

similar bodies in the Grenville Province, the

parental magma was probably derived from a


Table 2

Chemical analyses of metaigneous rocks from the northern Oaxacan Complex, southern Mexico

Sample Upper thrust slice Lower thrust slice

Charnockites Syenite Metagabbros OX-56 OX-57 OX-58 OX-59 OX-60 OX-61 OX-62 C-37 6398 6498 OC9810

OX-9 OX-10 OX-11 OX-15 OX-63 OX-64 OX-65 OX-66 OX-67 C-42 6098 OX-39 OX-41 OX-42 OX-43 66A98

SiO2 (wt.%) 62.48 60.87 54.62 60.05 63.11 63.12 48.34 49.00 60.00 53.02 59.6 43.53 45.93 44.73 46.71 48.9 37.80 35.42 58.38 46.16 46.68 62.11 56.37 53.10 36.4 36.9 59.4

TiO2 1.15 1.44 2.34 1.63 1.12 1.04 2.93 3.00 1.10 2.43 1.19 0.86 1.62 1.02 2.17 1.903 6.41 9.02 1.56 4.75 3.45 1.33 1.24 2.20 9.519 5.99 1.54

Al2O3 14.38 13.90 13.90 12.88 14.59 14.66 14.58 14.44 18.41 13.55 18.0 17.08 14.34 17.46 14.17 14.2 7.96 5.09 14.44 12.31 10.50 15.11 14.45 20.19 4.2 5.8 14.8

Fe2O3 7.49 9.32 11.69 10.63 7.32 6.87 14.10 13.00 3.99 12.12 4.6 12.91 13.04 11.84 14.40 14 20.40 26.58 12.04 15.46 24.18 6.71 11.46 5.36 26 36 12.3

MnO 0.11 0.13 0.15 0.16 0.11 0.10 0.22 0.21 0.06 0.17 0.066 0.28 0.26 0.24 0.34 0.277 0.30 0.40 0.25 0.22 0.40 0.09 0.19 0.07 0.373 0.906 0.223

MgO 1.07 3.86 2.33 1.49 1.02 0.96 5.15 5.00 0.95 3.01 1.14 10.16 6.94 8.68 5.85 5.3 4.99 7.19 1.15 3.14 1.91 1.37 2.51 1.59 7.4 1.1 0.696

CaO 3.31 0.00 5.39 4.07 3.23 3.14 8.36 8.08 2.32 5.89 2.45 8.94 10.42 9.15 10.14 11.5 11.08 8.51 4.11 7.80 5.76 2.67 3.27 7.20 9.7 6.4 3.6

Na2O 3.37 3.21 3.39 3.16 3.67 3.56 2.65 2.76 4.54 3.45 4.34 2.56 3.05 3.02 3.57 3.2 1.35 0.95 2.43 2.48 1.39 2.74 2.25 5.04 1.0 1.2 2.8

K2O 4.71 3.86 3.36 3.64 4.72 4.72 1.51 1.72 6.84 3.29 6.48 0.67 0.98 0.83 0.46 0.945 1.01 0.36 3.65 1.73 2.16 4.34 3.56 1.50 0.347 1.02 3.8

P2O5 0.37 0.49 0.95 0.56 0.37 0.35 0.85 1 0 1.21 0.341 0.08 0.24 0.1 0.27 0.278 5.85 3.89 1.22 2.77 2.12 1.07 0.92 1.38 4.7 2.91 0.822

L.O.I. 0.29 0.2 0.2 0.39 0.1 0.1 0.20 0 0 0.1 0.39 1.28 1.03 1.31 0.09 0.13 0.74 0.19 0.6 1.15 0.19 1.35 1.9 0.88 0 0 0.37

Totals 98.72 97.28 98.33 98.66 99.36 98.61 98.88 98.20 98.21 98.24 98.77 98.35 97.85 98.38 98.16 100.73 97.89 97.60 99.83 97.96 98.75 99.64 98.12 98.51 99.92 97.28 10035

Mg # 22.05 45.9 28.30 22.32 22.21 22.26 41.97 44.08 32.79 33.73 33.93 60.92 51.31 59.21 44.58 43.08 32.63 34.87 15.91 28.69 13.52 26.67 30.25 37.01 36.42 5.92 10.36

Cr (ppm) 18 6 8 18 �/ 13 37 36 10 2 5 83 103 66 63 58 2 �/ �/ 13 1 0 10 1 12 9 4

Ni �/ �/ �/ �/ �/ �/ �/ �/ 2 �/ 29 216 39 154 17 100 �/ �/ �/ �/ �/ �/ �/ 4 24 3 38

Co 6 18 25 14 10 12 51 37 �/ 27 47 89 81 71 54 89 44 60 6 31 34 18 13 18 73 60 98

Sc 22 19 9 18 12 18 10 15 17 5 6.32 11 6 4 16 39.16 6 19 11 11 25 18 24 8 �/ 1.56 31.72

V 87 120 219 118 62 75 337 315 48 226 38 280 715 282 788 400 72 243 10 118 36 63 46 48 217 17 20

Cu 7 12 14 11 6 6 21 20 4 19 41 14 43 56 87 218 37 34 16 36 30 8 21 10 47 46 51

Pb 20 17 18 15 19 20 6 6 19 16 18.26 7 3 4 4 3.54 5 9 5 8 12 6 6 1 5 3.03 9.26

Zn 120 141 175 167 104 104 154 146 58 181 76 111 123 212 180 162 262 363 117 274 431 69 177 76 596 540 229

Sn 4 3 �/ 2 �/ 3 �/ �/ 4 �/ �/ 2 2 �/ 3 �/ 1 �/ �/ 9 1 �/ 3 2 �/ �/ �/

Cs 0.4 �/ 0.1 �/ �/ �/ 0.1 �/ �/ �/ �/ 0.3 0.1 0.2 0.02 �/ 0.2 0.2 0.2 �/ 0.1 0.2 0.8 0.13 �/ �/ �/

Rb 112 112 87 89 105 105 22 41 70 80 45.33 17 7 16 3 7.33 7 3 30 31 25 35 33 16 9 9.61 47.61

Ba 2388 2036 2325 2266 2644 2737 625 581 2667 2306 2891 335 202 313 275 139 1285 181 4545 2206 3377 4482 4424 824 579 511 4313

Sr 385 345 615 336 393 393 485 473 515 711 458.74 319 202 340 286 277.12 714 439 501 1059 382 549 505 1483 494 128.11 440.99

Ga 18 17 16 15 17 19 14 15 19 14 �/ 13 14 14 13 �/ 2 �/ 32 12 �/ 25 17 21 �/ �/ �/

Ta 1.2 �/ 1.7 �/ �/ �/ 1.0 �/ �/ �/ 0.67 0.3 0.8 0.4 0.6 1.05 1.0 1.4 0.8 �/ 1.02 0.66 0.7 0.51 �/ 2 2.68

Nb 21.6 25 29.2 27 18 17 17.3 17 8 24 8.53 3 8.7 2.4 6.2 5.97 16.9 25.2 11.1 32 17.6 9.6 10.4 5.31 14 22.00 16.86

Hf 15.3 �/ 11.96 �/ �/ �/ 4.2 �/ �/ �/ 3.54 1.5 3.1 1.8 2.9 2.34 1.4 1.8 8.3 �/ 28.3 10.61 1.9 �/ �/ 17.66 10.84

Zr 636 711 593 884 603 567 338 320 695 499 376 66 133 79 145 130 141 168 1090 530 2968 1783 302 17.43 296 3658 1767

Y 50.8 67 85.7 62 44 44 63.4 50 19 66 19 23.8 53.9 18.6 42.0 32 90.1 66.2 74.1 58.0 85.4 29 26.8 14.73 122 79 21

Th 2.3 4.0 2.8 4.0 3.0 3.0 0.2 �/ 1 2 0.89 0.4 0.5 0.4 0.3 0.85 0.5 0.4 0.5 �/ 0.6 0.40 0.3 0.18 1 0.98 0.58

U 0.97 1 0.9 2 �/ 2 0.2 �/ 1 �/ 0.23 0.4 0.2 0.4 0.2 0.40 0.2 0.2 0.2 0.5 3.0 0.24 0.1 0.07 �/ 0.76 0.42

La 63.10 68 100 71 50 56 41.86 36 67 77 76.28 7.46 18.64 8.33 11.35 17.42 96.85 67.30 33.26 66.00 55.56 30.32 34.61 20.40 89 67.04 41.20

Ce 137.22 148 221 155 162 143 97.73 147 209 101 152.96 18.68 47.03 16.63 29.80 38.80 230 167.32 76.79 86 131.69 71.15 72.99 45.51 255 170.73 94.52

Pr 18.68 �/ �/ �/ �/ �/ 14.39 �/ �/ �/ 17.93 2.97 6.77 2.30 4.58 5.18 �/ �/ 11.98 �/ 20.17 10.77 10.67 6.55 34 25.85 13.44

Nd 75.63 65 99.10 50 49 43 63.75 41 58 74 70.49 14.07 28.56 10.82 21.24 22.43 �/ �/ 56.47 �/ �/ 48.81 50.26 21.85 132 125.41 61.27

Sm 14.46 �/ 24.94 �/ �/ �/ 13.84 �/ �/ �/ 10.10 3.61 7.36 2.80 6.08 5.68 30 25.90 12.09 �/ 20.18 9.88 9.57 5.85 29 25.82 12

Eu 4.26 �/ 4.61 �/ �/ �/ 3.63 �/ �/ �/ 5.34 1.23 2.03 1.19 1.83 1.64 �/ �/ �/ �/ �/ �/ �/ 2.51 8 6.54 10

Gd 12.44 �/ 20.66 �/ �/ �/ 12.92 �/ �/ �/ 7.29 3.83 8.30 3.28 6.87 5.49 30 22.69 12.33 �/ 19.34 7.85 8.47 5.10 28 23.16 10

Tb 1.66 �/ 2.70 �/ �/ �/ 1.94 �/ �/ �/ 0.93 0.61 1.42 0.52 1.15 1.06 3.74 2.68 1.86 �/ 2.53 1.03 1.03 0.58 3 3.62 2

Dy 9.90 �/ 16.08 �/ �/ �/ 11.27 �/ �/ �/ 4.45 3.94 9.46 3.16 7.30 6.95 18.60 13.67 12.31 �/ 14.95 5.39 5.13 2.91 19 21.23 9

Ho 1.97 �/ 3.24 �/ �/ �/ 2.33 �/ �/ �/ 0.67 0.86 2.06 0.69 1.58 1.50 3.36 2.46 2.87 �/ 3.16 1.12 0.94 0.52 3 4.29 1.64

Er 5.10 �/ 8.52 �/ �/ �/ 6.30 �/ �/ �/ 1.67 2.33 5.68 1.96 4.36 4.31 7.43 5.70 8.37 �/ 9.00 3.15 2.25 1.18 8 12.19 4

Tm 0.70 �/ 1.12 �/ �/ �/ 0.88 �/ �/ �/ 0.21 0.32 0.82 0.26 0.60 0.62 0.82 0.64 1.26 �/ 1.32 0.48 0.28 0.14 0.84 1.76 0.59

Yb 4.28 �/ 6.66 �/ �/ �/ 5.24 �/ �/ �/ 1.27 1.99 5.33 1.74 3.73 4.02 4.26 3.62 8.38 �/ 8.92 3.25 1.53 0.76 �/ 11.80 4

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mixed mantle-lower crustal source and was em-

placed in an extensional setting (Bourne, 1991).

5.1.2. Metagabbros

The metagabbros (OX-39-43) have SiO2 ranging

from 42 to 47 wt.% and a large spread of Mg#

from 0.60 to 0.45. They display a tholeiiticfractionation trend of increasing TiO2 (0.9�/2.2

wt.%) with Fe/Mg ratio accompanied by low Ti/V

ratios (B/30; Fig. 8). They are low in incompatible

trace elements (Table 2). Their REE patterns are

relatively flat (Fig. 9b) with Lan �/20�/60 and (La/

Yb)n ranging from 2 to 3. The flat LREE segment

of the pattern for these rocks probably reflects the

role of clinopyroxene. Their mantle-normalizedpatterns are concave with peaks at La and Ce and

display a variable depletion of high-field-strength

elements (Zr�/Nb�/Ta�/Ti; Fig. 10b). Their highly

variable concentrations of Mg, Cr, Ti and Zr are

consistent with field and petrographic observa-

tions that some of these rocks are probably, in

part, cumulates. They are tholeiitic gabbros similarto those found in the lower crust (Quick et al.,

1994; Sinigoi et al., 1994; Voshage et al., 1990).


The mafic-intermediate rocks (OX-55-62, C-36-

38) that are closely associated with anorthosite

show a significant range of chemical compositionsas exemplified by a variation of SiO2 from 35 to 62

wt.% and Mg# from 0.13 to 0.35. The rocks have

very high TiO2 (1.2�/9.0 wt.%), Fe2O3 (5.4�/27

wt.%) and P2O5 (0.9�/5.9 wt.%), reflecting presence

of Fe�/Ti oxides and apatite whereas a wide range

of Al2O3 (5�/20 wt.%) is due to significant varia-

tions of their plagioclase contents. Compared with

common basaltic rocks, CaO (2.5�/11 wt.%) andMgO (1.1�/7 wt.%) are low relative to Fe2O3

(Table 2). The abundances of Ba and Sr span a

wide range, but are typically high (up to 4400 and

1050 ppm, respectively). Concentrations of Rb are

usually low and, in conjunction with high Sr and

K, result in low Rb/Sr (0.03�/0.14) and high K/Rb

ratios (up to 800). Concentrations of Sc, Cr, Co

and Ni are all low (typically B/50 ppm) and inmany cases Cr and Ni are below the detection limit

(�/5 ppm). The abundance of V is also low,

mostly below 100 ppm, while Ti/V ratios are

uniformly high (�/100; Fig. 8), which distinguish

jotunites from most basaltic rocks (Owens et al.,

1993). The REE abundances are highly variable

(Fig. 9c). In general, LREE contents increase with

increasing P2O5. Correlation of P2O5 with LREEis probably due to accumulation of apatite. The

samples can be further subdivided into a group

with high Zr and a flat heavy REE pattern and the

second group with a sloping heavy REE (HREE)

pattern and lower Zr concentrations. The high

concentrations of Zr and flat HREE patterns

probably reflect an accumulation of zircon, in

addition to apatite and garnet. The samples withlowest total REE concentrations display a positive

Eu anomaly. The mantle-normalized trace element

patterns of the rocks of this unit show a distinct

enrichment of Ba relative to Rb and Th and an

enrichment of La and Ce relative to Nb�/Ta (Fig.

10c). The patterns of samples with flat HREE

Fig. 6. Zr/TiO2�/0.001 vs. SiO2 (wt.%) diagram of Winchester

and Floyd (1977) for the metaigneous rocks of the northern

Oaxacan Complex. Ab, alkali basalt; Sub-Ab, subalkaline

basalts; TrAn, trachyandesites. Squares, metagabbro in the

upper thrust slice; Circles, charnockite in the upper thrust slice;

Diamonds, jotunite (mafic-intermediate rocks in the lower

thrust slices).


display a positive Zr anomaly while the others

show no Zr anomaly or show a negative anomaly.

These rocks are closely comparable with those

of many mafic lithologies associated with anortho-

site massifs around the world (Owens et al., 1993;

Greenough and Owen, 1995; Geringer et al., 1998;

Icenhower et al., 1998), particularly oxide-apatite

gabbronorites and jotunites (mainly orthopyrox-

ene monzodiorites), in their mineral assemblages

(plagioclase�/clinopyroxene�/orthopyroxene�/

ilmenite�/magnetite�/apatite), their high Ti, Fe, P

and K but low Mg contents, and their high Ba, Sr,

LREE, Zr, Nb and low Cr and Ni. The low Mg#

as well as low Cr, and Ni abundances argue that

these rocks are not primary mantle melts. The

mineralogical and geochemical characteristics

have been explained as either evidence for partial

melting of mafic granulites of the lower crust or

extensive fractional crystallization of a mantle-

derived magma (Owens and Dymek, 1992). Our

data are consistent with the model invoked by

Icenhower et al. (1998) where the oxide-apatite

gabbronorites and jotunites were probably derived

from an enriched mantle source by partial melting

followed by a fractionation of olivine and ortho-

pyroxene. The fractional crystallization was ac-

companied by an increase of the concentrations of

P and Ti, which caused an extension of the

orthopyroxene stability field relative to that of

olivine. Fractional crystallization yielded the series

Fig. 7. Variations of K2O (wt.%), CaO (wt.%), CaO/Al2O3 (wt.%), and TiO2 (wt.%) relative to SiO2 for the metaigneous rocks of the

northern Oaxacan Complex.


of rocks ranging from anorthosites, leuconorites,

oxide-apatite gabbronorites to jotunites. Most

models for emplacement of such mafic magma

into continental crust are connected with incipient

rifting or mantle upwelling.

6. Origin of Oaxacan within-plate suites

The results presented in this paper indicate thatthere are two igneous suites of �/1157�/1130 and

1012 Ma ages in the northern Oaxacan Complex.

Although the chemistry of the older suite has both

within-plate and volcanic arc characteristics, the

enrichment in La and Ce relative to Nb, Ta, and

Th, is inferred to result from crustal contamina-

tion, a conclusion borne out by the negative oNd

values. The younger suite is closely associated with

anorthosite suggesting they form part of the same

igneous event, in which case they may be classified

as an anorthosite�/mangerite�/charnockite�/granite

(AMCG) suite. On the other hand, the Union

Zaragoza anorthosite in the upper thrust slice

could be associated with either intrusive event or

a third episode. If further geochronological work

establishes that the Unıon Zaragoza Anorthosite is

temporally associated with the older igneous suite,

then the rocks of the upper thrust slices may be

part of an older AMCG suite. The older suite is

presently located above the younger suite. This

may be a consequence of thrusting during the

granulite facies tectonothermal event dated at

Fig. 8. V vs. Ti/1000 diagram of Shervais (1982) for metaigneous rocks of the northern Oaxacan Complex. Squares, metagabbros in

the upper thrust slice; Circles, charnockite in the upper thrust slice; Diamonds, jotunite (mafic-intermediate rocks in the lower thrust

slices).


�/1004�/980 Ma or a subsequent deformational

event. On the other hand, the younger intrusive

suite could have been injected at lower structural

levels. There is also evidence of an earlier migma-

tization event dated at �/1100 Ma (Solari et al.,

2002). Taken together, these data indicate that

episodes of rift-related, within plate magmatism

immediately preceded tectonothermal episodes.

These data complement those from elsewhere in

Oaxaquia. In the southern Oaxacan Complex, an

arc complex was intruded by a �/1117 Ma, rift-

related granite followed by granulite facies meta-

morphism at 9889/5 Ma (Keppie et al., 2001). East

of the Oaxacan Complex in the Guichicovi Com-

plex (Fig. 1), the granulite facies metamorphism

occurred at 9869/4 Ma (Ruiz et al., 1999; Weber

and Kohler, 1999). In the Huiznopala Gneiss of

east�/central Mexico (Fig. 1), �/1200�/1150 Ma

arc magmatism was followed by intrusion of an

anorthosite/gabbro complex followed by granulite

Fig. 9. Chondrite-normalized rare-element patterns for the metaigneous rocks of the northern Oaxacan Complex: (a) charnockites of

the upper thrust slices; (b) metagabbros of the upper thrust slices; (c) mafic-intermediate rocks of the lower thrust slices. Normalizing

values after Sun (1982).


facies metamorphism at �/1000 Ma (Lawlor et al.,

1999). At Novillo, �/11759/16 Ma, arc or backarc

magmatism preceded intrusion of an 11329/34

Ma, AMCG suite followed by granulite faciesmetamorphism at �/980 Ma (Cameron et al.,

2002). Pb isotope data for feldspars separated

from igneous suites in all these inliers plot on a

tight linear array inferred to represent mixing

between two end-member sources that were ad-

jacent throughout the 1.2�/1.0 Ga period (Ca-

meron et al., 2002). These data indicate that the

Mexican Pb isotopes are more similar to theAndean massifs than the Grenville Province, and

suggests that the variation in the Pb isotopic

signatures may be related to other factors, such

as variations in the relative proportions of Ar-

chean and juvenile material contributing to the

source of the magmas. This conclusion does not

support the idea of dividing Oaxaquia along the

Trans-Mexican Volcanic Belt based upon empiri-cal correlation of whole-rock Pb isotopic signa-

tures (Ruiz et al., 1999). This together with the

synchroneity of events throughout these areas is

consistent with the concept of a single Oaxaquia

terrane (Ortega-Gutierrez et al., 1995). Further-

more, the main, �/1004�/980 Ma tectonothermal

event in Oaxaquia is apparently absent in Texas

which has a distinct Pb isotopic signature (Smith etal., 1997) suggesting that Oaxaquia is an exotic

terrane (Keppie and Ortega-Gutierrez, 1999).

Three extensional tectonic settings have gener-

ally been inferred for the AMCG and within-plate

magmatism in the Grenville Province: anorogenic

or incipient intercontinental rifting, intra- or back-

arc rifting, and convective thinning of the sub-

continental lithosphere following crustal thicken-ing (Windley, 1993; Corrigan and Hanmer, 1997;

Rivers, 1997). Arc magmatism in Oaxaquia may

have overlapped the emplacement of the �/1157�/

1130 Ma complex, and a rifted arc model is viable

for the older igneous suite. Although the location

of the �/1012 Ma AMCG complex within a

juvenile arc favors a rifted arc environment,

synchronous arc magmatism in Oaxaquia has notbeen recorded. On the other hand, the observation

that the younger �/1012 Ma Oaxacan AMCG,

within-plate complex immediately precedes an

orogenic event would appear to eliminate the third

model. However, convective thinning of the sub-

continental lithosphere can also occur where a

ridge collides with a trench, which can lead to

Fig. 10. Primitive mantle-normalized abundances of trace

elements in the metaigneous rocks of the northern Oaxacan

Complex: (a) charnockites of the upper thrust slices; (b)

metagabbros of the upper thrust slices; (c) mafic-intermediate

rocks of the lower thrust slices. Normalizing values after Sun

and McDonough (1989).


development of a slab window associated with

rifting and a switch from arc to rift magmatism

(Brown, 1998). It is also possible that synchronous

arc magmatism may exists in adjacent parts of

Rodinia. In order to evaluate these possibilities, we

have compiled a correlation chart to compare

AMCG and within-plate complexes, tectonother-

mal events and arc magmatism in potentially

correlative 1 Ga orogenic belts (Fig. 12).

Fig. 2 shows that most of the �/1 Ga AMCG/

within-plate complexes are located in the Grenville

Province of eastern Laurentia, the Sveconorwe-

gian Orogen of southern Baltica and in the

Andean massifs, such as the Santa Marta massif

and the Merida Andes. Their absence is notable in

the Grenville Orogen of Texas (Mosher, 1998),

Cuyania (a terrane presently in western Argentina

that may have originated in the Ouachita Embay-

ment: Thomas and Astini, 1996), in the Arequipa

massif of Peru (Wasteneys et al., 1995), and in the

Sunsas Orogen of Bolivia (Litherland et al., 1986).

They are rarely reported in Antarctica (Tingey,

1991). The two AMCG/within-plate suites in the

Oaxacan Complex appear to be synchronous with

1170�/1120 and 1025�/1010 Ma suites in the

Grenville Province, however, the age of the

AMCG suite(s) in the Andean massifs in currently

unknown (Restrepo-Pace et al., 1997) making

temporal comparisons impossible. Arc magmatism

appears to have ceased throughout the adjacent �/

1 Ga orogens by �/1150 Ma, with an exception in

the Sveconorwegian orogen where syntectonic,

calcalkaline magmatism has been dated at �/

1040 Ma (Bingen et al., 1993), and �/920 Ma

arc magmatism in the Oaxacan Complex (Ortega-

Obregon, 2002). On the other hand, the Avalonian

basement appears to be made up of a juvenile �/

1.2�/1.0 Ga arc (Murphy et al., 2001), and so could

be representative of arc activity adjacent to the �/

1012 Ma AMCG/within-plate complexes in Oax-

aquia and the Grenville Province. Thus, several

tectonic settings appear to be possible for the �/

1012 Ma AMCG suite in Oaxaquia: an intra-arc

rift, a backarc rift, or a slab window rift. These

environments would allow access of hot mantle to

the base of the crust, which would induce partial

melting and AMCG/within-plate magma produc-

tion. All of these environments require that

Oaxaquia be placed on the periphery of Rodinia

with a subducting ocean on one side. Whereas

most �/1 Ga orogens are dominated by remobi-

lized basement, Oaxaquia appears to be a juvenile

�/1.4�/1.0 Ma composite arc terrane, a feature

mainly found in the outboard Grenvillian terranes,

the allochthonous Appalachian terranes, and the

Andean terranes (Fig. 12). Based upon faunal

affinities in Lower Paleozoic rocks in Oaxaquia

(Robison and Pantoja-Alor, 1968; Robison, writ-

Fig. 11. Variations of (a) Nb vs. Y; and (b) Rb vs. (Y�/Nb) in felsic metaigneous rocks of the northern Oaxacan Complex. Fields are

after Pearce et al. (1984): VAG, volcanic arc granites; syn-COLG, syn-collisional granites; WPG, within-plate granite; ORG, ocean

ridge granite.


ten communication 1998; Boucot et al., 1997),

Keppie et al. (2001) favored a provenance for

Oaxaquia off northern Amazonia (location #2 in

Fig. 2). Given the lack of tectonic events in

Oaxaquia between the �/1004�/980 Ma Zapotecan

Orogeny and the Lower Paleozoic, such a prove-

nance may reasonably be extended back to �/1

Ga.

However, recent work suggests that the juxta-

position of eastern Laurentia and Arequipa�/

Amazonia may not be valid. Thus, Ramos and

Aleman (2000) show a Brasiliano orogenic belt

between the Peruvian Arequipa massif and the

Amazon craton (Fig. 2 inset), which implies that

they cannot have been juxtaposed until the Neo-

proterozoic. Furthermore, Loewy et al. (2000)

have shown that the correlation between the

Scottish Dalradian and the Peruvian Arequipa

massif proposed by Dalziel (1994) is not supported

by recent data (Fig. 1 inset). These results allow

several options for the provenance of Oaxaquia

and the Chortis block. They may be placed any-

where around the northern, western and southern

margins of Amazonia with an open ocean farther

Fig. 12. Correlation chart for events between 1200 and 900 Ma in Oaxaquia and potential correlatives. References for data: Baltica

(Larson, 2000, and references therein); Grenville (Rivers, 1997; Aleinikoff et al., 2000, and references therein); Blair River (Miller and

Barr, 2000); Goochland (Aleinikoff et al., 1996); Texas (Mosher, 1998, and references therein); Oaxaquia (this paper and Keppie et al.,

2001, and references therein); Cuyania (Baldo et al., 1997, and references therein); northern Andean massifs (Aleman and Ramos,

2000, and references therein); Arequipa (Wasteneys et al., 1995); Sunsas (Tassinari et al., 2000, and references therein); Musgrave

(White et al., 1999, and references therein); Coats Maud Land, Antarctica (Tingey, 1991, and references therein); Natal-Namaqua

(Jacobs et al., 1993; Thomas et al., 1993, and references therein); East Antarctica (Tingey, 1991, and references therein).


outboard (c.f. Keppie and Ortega-Gutierrez, 1999;

Keppie and Ramos, 1999; Ramos and Aleman,

2000). They could be placed adjacent to the

Arequipa massif as a microcontinental-juvenile

arc terrane within a Grenville ocean (Location

#3 in Fig. 2; e.g. Keppie and Ortega-Gutierrez,

1999). They could also be placed outboard of the

eastern margin of Laurentia and Baltica (Location

#3 in Fig. 2; e.g. Cawood et al., 2001).

The �/980�/1004 Ma Zapotecan event correlates

both temporally and in P �/T conditions with that

in the Andean massifs (Arequipa, Santa Marta

and Merida Andes: Wasteneys et al., 1995; Re-

strepo-Pace et al., 1997), with the allochthonous

Appalachian massifs (Goochland and Blair River

terranes: Aleinikoff et al., 1996; Miller and Barr,

2000), and with the 1060�/1000 Ma Sveconorwe-

gian orogeny in Scandinavia (Romer and Smeds,

1996; Romer, 1996). Although it is synchronous

with the Rigolet event in the Grenville Province,

the Rigolet effects are limited to the margin of

the orogen (Rivers, 1997). Current data are

insufficient to resolve these alternative options

for the provenance of Oaxaquia. More data

from Oaxaquia and northern South America

are required to provide better constraints on

correlations.

Acknowledgements

We would like to thank Dr Joaquin Ruiz and Dr

Pedro Restrepo-Pace for their constructive reviews

of the manuscript. Funding for various aspects of

this project were provided by CONACyT grants

(0255P-T9506 and 25705-T), PAPIIT grants

(IN116999 and IN10799) to J.D. Keppie and F.

Ortega-Gutierrez, NSERC grant to J. Dostal, a

NSF Grant EAR 9909459 to K.L. Cameron, and

MEXUS grant to K.L. Cameron and F. Ortega-

Gutierrez. We would like to thank Pete Holden for

assistance with analyses, Elena Centeno-Garcia

for assistance with sample collection, Carlos

Ortega for sample preparation, and Jose Luis

Arce for drafting some of the figures.

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Geochronology and geochemistry of Grenvillian igneous suites in the northern Oaxacan Complex, southern Mexico: tectonic implications

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