990 and 1100 Ma Grenvillian tectonothermal events in the northern Oaxacan Complex, southern Mexico:...

990 and 1100 Ma Grenvillian tectonothermal events in the northern

Oaxacan Complex, southern Mexico: roots of an orogen

L.A. Solaria,*, J.D. Keppiea, F. Ortega-Gutierreza, K.L. Cameronb,R. Lopezc, W.E. Hamesd

a Instituto de Geologıa, Universidad Nacional Autonoma de Mexico, 04510 Mexico, D.F., MexicobEarth Sciences Department, University of California, Santa Cruz, CA 95064, USA

cGeology Department, West Valley College, Saratoga, CA 95070, USAdDepartment of Geology, Auburn University, Auburn, AL 36830, USA

Received 14 February 2002; accepted 13 June 2002

Abstract

Inliers of f 1.0–1.3 Ga rocks occur throughout Mexico and form the basement of the Oaxaquia microcontinent. In the

northern part of the largest inlier in southern Mexico, rocks of the Oaxacan Complex consist of the following structural

sequence of units (from bottom to top), which protolith ages are: (1) Huitzo unit: a 1012F 12 Ma anorthosite–mangerite–

charnockite–granite (AMCG) suite; (2) El Catrın unit: z 1350 Ma orthogneiss migmatized at 1106F 6 Ma; and (3) El

Marquez unit: z 1140 Ma para- and orthogneisses. These rocks were affected by two major tectonothermal events that are

dated using U–Pb isotopic analyses of zircon: (a) the 1106F 6 Ma Olmecan event produced a migmatitic or metamorphic

differentiation banding folded by isoclinal folds; and (b) the 1004–978F 3 Ma Zapotecan event produced at least two sets of

structures: (Z1) recumbent, isoclinal, Class 1C/3 folds with gently NW-plunging fold axes that are parallel to mineral and

stretched quartz lineations under granulite facies metamorphism; and (Z2) tight, upright, subhorizontal WNW- to NNE-trending

folds accompanied by development of brown hornblende at upper amphibolite facies metamorphic conditions. Cooling through

500 jC at 977F 12 Ma is documented by 40Ar/39Ar analyses of hornblende. Fold mechanisms operating in the northern

Oaxacan Complex under Zapotecan granulite facies metamorphism include flexural and tangential– longitudinal strain

accompanied by intense flattening and stretching parallel to the fold axes. Subsequent Phanerozoic deformation includes

thrusting and upright folding under lower-grade metamorphic conditions. The Zapotecan event is widespread throughout

Oaxaquia, and took crustal rocks to a depth of f 25–30 km by orogenic crustal thickening, and is here designated as

Zapotecan Orogeny. Modern analogues for Zapotecan granulite facies metamorphism and deformation occur in middle to lower

crustal portion of subduction and collisional orogens. Contemporaneous tectonothermal events took place throughout Oaxaquia,

and in various parts of the Genvillian orogen in Laurentia and Amazonia.

D 2003 Elsevier Science B.V. All rights reserved.

Keywords: Oaxaquia; Mexico; Grenville; U–Pb geochronology; Granulite metamorphism

0040-1951/03/$ - see front matter D 2003 Elsevier Science B.V. All rights reserved.

doi:10.1016/S0040-1951(03)00025-8

* Corresponding author. Tel.: +52-55-5622-4263x113; fax: +52-55-5622-4289.

E-mail address: [email protected] (L.A. Solari).

www.elsevier.com/locate/tecto

Tectonophysics 365 (2003) 257–282

1. Introduction

The Oaxacan Complex (OC) is the largest inlier

of f 1 Ga rocks in Mexico and underlies more than

10,000 km2 of southern Mexico (Fig. 1A). Other

similar smaller inliers, borehole samples and xeno-

liths in Tertiary lavas led Ortega-Gutierrez et al.

(1995) to conclude that the backbone of Mexico is

underlain by a similar f 1.0 Ga basement, which

they named Oaxaquia (Fig. 1A). It may continue into

the Chortis Block of Guatemala and Honduras

(Donnelly et al., 1990; Manton, 1996; Nelson et

al., 1997), which is inferred to have lain along the

Pacific margin of Mexico before being sinistrally

displaced in the Oligocene (Schaaf et al., 1995). As

such, Oaxaquia and the Chortis block constitute

major pieces that need to be considered in recon-

structing f 1 Ga Rodinia. However, their inferred

location in recent reconstructions of Rodinia differs

widely. For example, Karlstrom et al. (1999) and

Burrett and Berry (2000) infer that Oaxaquia is the

southern extension of the Laurentian Grenville Oro-

gen, and represents a connecting segment to the

f 1.0 Ga orogens of eastern Australia. On the other

hand, other authors have proposed that Oaxaquia is

an exotic terrane of either Amazonian or northeastern

Laurentian provenance (Keppie, 1977; Ballard et al.,

1989; Yanez et al., 1991; Keppie and Ortega-Gutier-

rez, 1995, 1999; Ortega-Gutierrez et al., 1999). The

presence of Ordovician and Silurian fauna with

Gondwanan affinity in rocks unconformably above

the 1 Ga rocks in Mexico (Robinson and Pantoja-

Alor, 1968; Boucot et al., 1997) is consistent with an

Amazonian location during the lower Paleozoic. As

no Neoproterozoic tectonic event has been recorded

in Oaxaquia, it is probably fair to assume that this

same location also applies to its f 1 Ga position.

Ruız et al. (1999) divided Oaxaquia into two pieces

along the Trans-Mexican Volcanic Belt (Fig. 1A),

and suggested that the northern piece was continuous

with southern Laurentia, and the southern one was

derived from Gondwana. In this paper, we document

that the northern Oaxacan Complex has undergone a

complex structural and metamorphic history that

involved two high-grade tectonothermal events at

f 1100 and f 990 Ma (named Olmecan and Zapo-

tecan, respectively). These are tentatively interpreted

as extensional and either Andean, arc–continent or

continent–continent collision events, respectively,

which occurred along the margin of Amazonia.

2. Geological setting

The Oaxacan Complex consists of para- and

orthogneisses, the latter having protolith U–Pb zircon

ages ranging from z 1134 to 1150 and f 1012 Ma

(Keppie et al., 2003), which were affected by granulite

facies metamorphism (700–750 jC and 7.2–8.2 kb,

Mora et al., 1986) at about 990 Ma (Keppie et al.,

2001, 2003). Previous structural studies of the Oax-

acan Complex by Kesler and Heath (1970) and Kesler

(1973) conclude that it was only affected by one phase

of deformation.

In several other areas in Mexico, f 1.0 Ga rocks

are also exposed (Fig. 1A): near Ciudad Victoria

(Novillo Gneiss, Silver et al., 1994), at Molango

(Huiznopala Gneiss, Lawlor et al., 1999), and around

La Mixtequita in eastern Oaxaca State (Guichicovi

Complex, Weber and Kohler, 1999; Ruız et al., 1999).

Published U–Pb igneous protolith ages are f 1000

and 1150–1200 Ma in Huiznopala (Lawlor et al.,

1999) and 1231F 43 Ma (upper intercept) in the

Guichicovi Complex (Weber and Kohler, 1999).

Granulite facies metamorphism was also dated by

U–Pb at f 1000 Ma in Huiznopala (Lawlor et al.,

1999), at f 986 Ma (Ruız et al., 1999) and 975F 36

Ma (Weber and Kohler, 1999) in the Guichicovi

Complex, and is poorly constrained between 928F18 Ma (K–Ar on phlogopite, Denison et al., 1971)

and 1018F 3 Ma (U–Pb on concordant zircon, Silver

et al., 1994) in the Novillo Gneiss.

Available major, trace and REE geochemistry, as

well as Nd and common Pb isotopic data suggest that

these f 1.0–1.2 Ga igneous protoliths constitute a

magmatic arc intruded by two series of intraplate and

AMCG (anorthosite–mangerite–charnockite–gran-

ite) suites (Patchett and Ruız, 1987; Ruız et al.,

1988, 1999; Lawlor et al., 1999; Weber and Kohler,

1999; Lopez et al., 2001; Keppie et al., 2001a,b). The

absence of a Paleoproterozoic or Archean rocks

beneath Oaxaquia is suggested by the fact that no

U–Pb upper intercept ages older than f 1400 Ma

have been recorded and Nd model ages range

between f 1.3 and f 1.6 Ga (Patchett and Ruız,

1987; Lawlor et al., 1999).

L.A. Solari et al. / Tectonophysics 365 (2003) 257–282258

Fig. 1. (A) Oaxaquia and the main outcrops of Grenvillian-age rocks in Mexico. The box shows the location of the area studied in this work; TMVB is the Trans-Mexican Volcanic

Belt. P–M fault is the Polochic–Motagua system, SE of which crops out the Chortis block. (B) Geological map of the studied area, northern Oaxacan Complex. (C) Reconstructed

structural column for the studied area. Both figures contain U–Pb and Ar–Ar sample locations.

L.A.Solariet

al./Tecto

nophysics

365(2003)257–282

259

3. Lithologies of structural units

Mapping of the northern Oaxacan Complex reveals

that it consists of a series of thrust slices composed of

(from bottom to top) (Fig. 1B,C): (1) the Huitzo

structural unit; (2) the El Catrın migmatite structural

unit; and (3) the El Marquez structural unit.

The Huitzo structural unit is made up of a f 3-km-

thick AMCG suite consisting of meta-anorthosite and

intercalated Fe-metadiorite, metagabbro, mafic cumu-

lates and garnet-bearing charnockite with protolith

ages of 1012F 12 Ma (Keppie et al., 2003). The Fe-

metadiorites and anorthosites are intimately inter-

leaved as a result of intense deformation that reor-

iented primary, magmatic contacts parallel to the

foliation. The Fe-metadiorites are dark grey in colour,

and vary from massive to intensely lineated in shear

zones. Granulite facies metamorphism is indicated by

mineralogy composed of plagioclase (An 30–45%),

orthopyroxene, augitic clinopyroxene, brown horn-

blende, rare titaniferous biotite, magnetite and ilmen-

ite, with abundant accessory apatite, and granoblastic

textures, with the typical triple junctions at 120j in themore massive units. Clinopyroxene is replaced by

amphibole ranging from tremolite–actinolite to antho-

phyllite pseudomorphs, whereas magnetite and ilmen-

ite are often surrounded by coronas of garnet, which

are typical of isobaric cooling conditions (Harley,

1989). Low-grade, late coronas of actinolite, biotite

and/or epidote indicate hydration and retrograde meta-

morphism under upper greenschist to amphibolite

facies. The metagabbros are dark green to black, and

composed of augite, plagioclase, brown to green

hornblende, rare biotite and relict hypersthene. Pres-

ence of this last mineral, together with textural

relationships between hornblende, biotite and clino-

pyroxene, suggest these metagabbros also underwent

granulite metamorphism. The anorthosite is white

and massive in the middle of the pluton, becoming

more foliated towards its margins. It is made up of

plagioclase (An 20–30%), magnetite and ilme-

niteFK-feldspar and quartz, and secondary calcite

and epidote. Nelsonites in the anorthosites are com-

posed of apatite, ilmenite and minor amounts of

magnetite. In the Adirondack Highlands of the

Laurentian Grenville Province, the nelsonites are

considered to be immiscible liquids, comagmatic

with the anorthosites (Darling and Florence, 1995).

The anorthosite–gabbro–Fe–metadiorite is in tec-

tonic contact with the overlying dark green, weakly

foliated mafic gneiss and garnet-bearing charnockite.

The tectonic contact is characterized by a 15-m-thick

greenschist to lower amphibolite shear zone (Fig.

1B,C) with top-to-the-SE kinematics (Solari, 2001).

The mafic gneisses are composed of plagioclase (An

25%), hypersthene, augite, hornblende, magnetite,

garnet, zircon and minor amount of alkalic feldspar

and quartz, which indicate that they also underwent

granulite facies metamorphism. One of the studied

samples belonging to this unit contains magmatic

pigeonite, with hypersthene exolutions. The chemistry

of these mafic gneisses suggests that they represent

cumulates (Keppie et al., 2003). A 200-m-thick slice

of pale gray, foliated charnockite is intercalated with

these gneisses. Its contacts are parallel to the banding,

and do not show any shearing. The charnockite is

composed of perthitic feldspar, quartz, augite, hyper-

sthene, garnet, plagioclase (An 30%) and accessory

apatite and zircon, an assemblage indicative of gran-

ulite facies metamorphism.

The El Catrın Migmatite structural unit comprises

migmatitic gneisses (about 2-km thick) that lie struc-

turally above the charnockites and the mafic gneisses.

Field relationships show that the mafic gneisses and

anorthosites of the Huitzo structural unit intruded these

migmatites (Solari, 2001, p.26). Migmatite in the type

locality is represented by a stromatic to nebulitic, light

grey leucosome in a dark gray to bluish mesosome.

The leucosome consists of quartz, alkali feldspar,

plagioclase (An 20–45%), secondary epidote, calcite

and chloritized biotite. The mesosome, in addition to

those minerals characterizing the leucosome, contains

augitic clinopyroxene, hypersthene, rare hornblende

and, in few samples, scapolite. This mineral assem-

blage was produced by granulite facies metamor-

phism, whereas secondary minerals were produced

by retrograde metamorphism. The mesosome, leuco-

some and the migmatitic fabric are only preserved in

the middle of a low-strain window. Outside this, the

migmatitic gneisses grade into striped gneiss, in which

subsequent deformation is stronger and generally

associated with a granulite metamorphic foliation

defined by augite and hypersthene. Although geo-

chemical data are not available, the petrography sug-

gests that the protolith of the migmatite was gabbroic–

dioritic in composition.


Intercalated paragneiss and orthogneiss of the El

Marquez structural unit immediately overlie the mig-

matite (Fig. 1B,C). The contact between the two is

characterized by a 150-m-thick biotite- and musco-

vite-bearing paragneiss, mylonitized under greenschist

facies, with a top-to-the-E sense of shear (Solari,

2001). The paragneisses are represented by a variety

of lithologies ranging from quartz–feldspar–garnet

(pyrope-almandine) gneiss, through two pyroxene–

quartz–feldspar gneiss, and mica–graphite–silliman-

ite–rutile gneiss to marbles and calcsilicates. Textures

are generally granoblastic. The impure marbles are

generally composed of calcite with abundant acces-

sory minerals such as diopside, olivine, purple fas-

saite, graphite, quartz, wollastonite and phlogopite.

The calcsilicates are composed of scapolite, micro-

cline, diopside, titanite, graphite, phlogopite and ac-

cessory calcite. Although some of these marbles and

calcsilicates are foliated and concordant with the

regional banding, others cut across the structures

and are massive. Pinnitized cordierite and sapphirine

rarely occur in the northern OC paragneisses. These

different parageneses and the textures indicate that the

paragneisses also underwent granulite facies meta-

morphism. Based upon the abundance of sulphur-

and chlorine-rich scapolite, alkali feldsparF gypsum

and anhydrite within this structural unit, Ortega-

Gutierrez (1984) proposed that the calcsilicates were

originally evaporites interbedded with carbonates,

arkoses, felsic igneous rocks (now quartzo-feldspathic

gneiss), marls, alkali basalt/dolerite (now amphibolite)

and magnesian clays (now ultramafic rocks). He

inferred that this association was deposited in a

continental rift environment. Minor igneous bodies

such as amphibolites and pegmatites, as well as major

charnockitic, meta-syenitic and meta gabbroic bands,

up to 500-m thick, intrude these paragneisses. The

charnockites are characterized by quartz, perthitic

feldspar, plagioclase (An 35%), hypersthene, augite

and hornblende. Metasyenite is mainly composed of

abundant meso-perthitic feldspar, quartz, hornblende,

biotite, clinopyroxene and relict hypersthene, whereas

meta-gabbros are composed of augite, plagioclase,

hornblende and magnetite. Zircon, Fe–Ti ore and

apatite are particularly abundant accessory minerals.

Their parageneses indicate that these rocks were

affected by granulite facies metamorphism that over-

printed the original, magmatic fabric. Previous U–Pb

zircon analyses carried on metasyenite, charnockite

and metagabbros indicate their crystallization ages

range between f 1134 and f 1230 Ma (Keppie et

al., 2003).

4. Structures and age constraints

The tectonic history of the three main structural

units of the northern Oaxacan Complex is as follows:

(i) in the f 1012MaHuitzo unit, two sets of structures

developed under granulite facies metamorphic condi-

tions; (ii) in the El Catrın unit, two sets of structures

associated with migmatization are preserved in a strain

window, and these are overprinted by two sets of struc-

tures formed under granulite and amphibolite facies

metamorphic conditions, respectively; and (iii) the

z 1134 Ma El Marquez unit records four sets of

structures, the first three of which are associated with

granulite facies minerals, whereas the last set devel-

oped under amphibolite facies conditions. On the basis

of age constraints (see below), the granulite–upper

amphibolite facies structures affecting the Huitzo unit

and their correlatives in the other units are assigned to

the Zapotecan tectonothermal event dated at f 978–

1004F 3 Ma, whereas the structures associated with

migmatization in the El Catrın unit are assigned to the

Olmecan tectonothermal event, dated at F 1100 Ma.

These names are derived from the pre-Hispanic cul-

tures that inhabited the Oaxacan region. Structures

associated with these two tectonothermal events are

overprinted by other structures developed at lower

amphibolite–subgreenschist facies metamorphic con-

ditions, such as thrusts, shear zones, and upright, NW-

trending open folds, which are Phanerozoic in age

(Solari, 2001). This paper is limited to the Zapotecan

and Olmecan events, which are designated by Z and u

subscripts, respectively, to constrain planar (S), and

linear (L) structures, and folds (F). In each structural

unit, these events are further distinguished by the

subscripts h (Huitzo structural unit), c (El Catrın

structural unit), or m (El Marquez structural unit). A

justification for the correlation between the structures

observed in each of these three structural unit and their

assignment to Olmecan or Zapotecan will be proposed

in the Summary section, but rather than introduce two

numbering schemes, the conclusions are used through-

out the paper.

L.A. Solari et al. / Tectonophysics 365 (2003) 257–282 261

Table

1

U–Pbgeochronologyfortheselected

samples,northernOaxacan

Complex,southernMexico

Fraction

Weight

(mg)

U (ppm)

TotalPb

(ppm)

Com.

Pb(pg)

206Pb/204Pb

206Pba/238U

206Pba/238U

(%err.)

207Pba/235U

207Pba/235U

(%err.)

207Pba/206Pba

207Pba/206Pba

(%err.)

206Pba/238U

207Pba/235U

207Pba/206Pba

Disc.

(%)

Raw

datab

Atomic

ratiosc

Age(M

a)d

ElCatrın

migmatite,neosome,

sample

67B98

(1)sng,elong,abr

0.015

78

18

22

695

0.19890

0.25

2.1821

0.32

0.07956

0.19

1169

1175

1186F4

1.4

(2)sng,rnd,abr

0.029

67

12

92473

0.17978

0.14

1.8613

0.19

0.07509

0.13

1066

1067

1071F3

0.5

(3)eq,abr,11grn

0.125

67

12

811,602

0.18524

0.09

1.9586

0.10

0.07668

0.04

1096

1101

1113F1

1.6

(4)2grns,irg

0.012

91

16

10

11,539

0.17607

0.10

1.8084

0.11

0.07449

0.05

1045

1048

1055F1

0.9

(5)eq,abr,sng

0.052

68

14

18

475

0.18387

0.24

1.9276

0.31

0.07603

0.20

1088

1091

1096F4

0.7

(6)pnk,sng,abr,cir

0.015

527

112

31

3135

0.19536

0.25

2.1265

0.26

0.07895

0.06

1150

1157

1171F2

1.8

(7)abr,sng,tip

0.020

26

54

1507

0.18726

0.58

1.9733

0.65

0.07643

0.28

1107

1106

1106F6

0.0

(8)CL,plkd,abr,sng

0.040

84

18

25

1663

0.19438

0.39

2.1044

0.42

0.07852

0.15

1145

1150

1160F3

1.3

(9)sng,pnk,abr,eq

0.010

262

57

21

1511

0.19633

0.19

2.1541

0.42

0.07957

0.35

1156

1166

1187F7

2.6

ElCatrın

migmatite,paleosome,

sample

67A98

(10)sng,prsm,fr

0.009

688

154

31

2784

0.21850

0.25

2.5355

0.28

0.08416

0.12

1274

1282

1296F2

1.7

(11)4grns,elong,pnk,crk

0.013

524

120

22

4278

0.21689

0.51

2.5059

0.54

0.08380

0.18

1265

1274

1288F4

1.8

(12)lbc,

abr,sng

0.013

650

138

21

5300

0.21050

0.10

2.3793

0.11

0.08198

0.04

1231

1236

1245F1

1.1

(13)eq,rnd,abr,sng

0.008

178

52

20

556

0.18044

0.60

1.9009

0.69

0.07640

0.34

1069

1081

1106F7

3.3

(14)eq,rnd,abr,sng

0.007

745

156

89108

0.21084

0.13

2.4214

0.13

0.08329

0.04

1233

1249

1276F1

3.3

(15)irg,abr,sng

0.008

818

179

10

9107

0.21567

0.09

2.4859

0.10

0.08360

0.04

1259

1268

1283F1

1.9

(16)irg,abr,2grn

0.008

649

139

88440

0.20859

0.09

2.3659

0.10

0.08226

0.04

1221

1232

1252F1

2.5

Folded

pegmatite

ElCuajilote,sample

6898

(17)elong,abr,sng

0.052

41

89

2097

0.18846

0.35

2.0025

0.38

0.07706

0.15

1113

1116

1123F3

0.9

(18)elong,abr,5grn

0.060

45

93

5546

0.18468

0.28

1.9584

0.31

0.07691

0.14

1092

1101

1119F3

2.4

Sem

ipelitic

metasediment,sample

7098

(19)eq,abr,sng

0.021

683

111

21

6611

0.16672

0.39

1.6613

0.41

0.07227

0.12

994

994

994F3

0.0

(20)abr,sph,multif,2grn

0.014

220

39

93859

0.18016

0.20

1.8807

0.25

0.07571

0.16

1068

1074

1087F3

1.8

(21)abr,sph,multif,sng

0.010

283

51

14

2364

0.18392

0.34

1.9460

0.40

0.07674

0.21

1088

1097

1114F4

2.3

Pre-tectonic

pegmatite,sample

66B98

(22)abr,lbc,

sng

0.047

96

20

38

1416

0.19083

0.11

2.0287

0.14

0.07710

0.08

1126

1125

1124F2

�0.2

(23)abr,fr,3grn

0.071

92

19

117155

0.19425

0.16

2.0857

0.18

0.07787

0.07

1144

1144

1144F2

0.0

(24)abr,fr,2grn

0.051

106

22

116451

0.19582

0.34

2.1069

0.37

0.07804

0.14

1153

1151

1148F3

�0.4

(25)abr,lbc,

sng

0.035

147

31

1118,544

0.19183

0.07

2.0566

0.08

0.07776

0.04

1131

1134

1141F2

0.9

(26)abr,sng,rndto

prsm

0.158

124

24

101

2143

0.17928

0.15

1.8582

0.16

0.07517

0.05

1063

1066

1073F2

0.9

(27)abr,fr,sng

0.031

71

14

16

1654

0.18484

0.39

1.9535

0.43

0.07665

0.19

1093

1100

1112F4

1.7

Flattened

axialplanarpegmatite,sample

66C98

(28)lbc,

flat,abr,mg

0.025

724

122

1114,104

0.16213

0.21

1.6000

0.22

0.07157

0.07

969

970

974F2

0.6

(29)lbc,

mg

0.112

563

96

25

15,261

0.16258

0.18

1.6065

0.19

0.07167

0.06

971

973

977F2

0.6

(30)stby,

srp.edges,clear,7grn

0.086

450

81

100

2949

0.16370

0.19

1.6188

0.25

0.07172

0.16

977

978

978F3

0.1

(31)lbc,

abr,mg

0.124

406

75

32

8535

0.16416

0.18

1.6278

0.21

0.07192

0.09

980

981

984F2

0.4

(32)stby,

srp.edges,abr,mg

0.028

912

149

55

4566

0.15984

0.13

1.5750

0.14

0.07147

0.04

956

960

971F2

1.6

Late

tectonic

pegmatite,sample

66D98

(33)abr,sng,stubbyto

rnd

0.082

288

50

19

12,196

0.16339

0.17

1.6145

0.17

0.07167

0.04

976

976

976F1

0.0

(34)abr,sng,stubbyto

rnd

0.045

937

168

22

19,279

0.16263

0.18

1.6081

0.19

0.07171

0.05

971

973

978F1

0.7

(35)abr,sng,elg

0.032

498

88

29

5530

0.16349

0.60

1.6164

0.62

0.07171

0.16

976

977

978F3

0.2

(36)abr,sng,elg

0.032

365

65

37

3202

0.16403

0.25

1.6222

0.26

0.07173

0.07

979

979

978F2

�0.1

ElTecolote

pegmatite,sample

6998

(37)fr,abr

0.188

123

20

23

9933

0.16241

0.20

1.6032

0.23

0.07159

0.12

970

971

975F1

0.5

(38)fr,abr

0.019

159

28

414,039

0.16434

0.44

1.6256

0.44

0.07174

0.04

980

980

979F3

�0.1


Several samples were collected for zircon U–Pb

isotopic analysis to bracket the age of the observed

tectonothermal events: two samples of migmatite

(67A98 paleosome and 67B98 neosome) of the El

Catrın unit, three pegmatites from one locality in the

El Marquez unit (66B98, 66C98, 66D98 that are

early, middle and late with respect to the local

fabrics), a pegmatite in the El Catrın unit that is

folded and sheared together with its host rock (6898),

a late-tectonic, high-grade pegmatite as well as a

metasediment (7098), both belonging to El Marquez

unit (6998), and finally a hornblende-bearing meta-

gabbro (OC9901) in the Huitzo unit for 40Ar/39Ar

isotopic analysis (Fig. 1b,c). Zircon separation,

chemistry and mass spectrometry were performed at

University of California, Santa Cruz (UCSC) follow-

ing the analytical procedures described in Lopez et

al. (2001). Isotopic ratios were reduced, and the

errors assessed, using the program PbDat (Ludwig,

1991) (Table 1) and the concordia diagrams plotted

using Isoplot for Excel v. 2.49 (Ludwig, 2001). The40Ar/39Ar laser fusion analyses of hand-picked horn-

blende crystals (Table 2) were performed at the

Massachusetts Institute of Technology using the

methodology outlined in Hames and Bowring

(1994).

Table 240Ar–39Ar geochronology for sample OC 9901, northern Oaxacan Complex, southern Mexico

Analysis 39Ar/40Ar

(� 10� 2)

36Ar/40Ar

(� 10� 5)

38Ar/40Ar

(� 10� 4)

37Ar/40Ar K/Ca K/Cl Ca/Cl %40Ar* Agesa

1 1.48F 0.0093 2.00F 1.06 2.31F 0.0630 0.0401F 0.0003 0.192 6.8 35.5 99.4 1021.9F 6.4

2 1.56F 0.0094 1.60F 0.984 2.06F 0.0672 0.0403F 0.0002 0.201 8.1 40 99.5 981.3F 5.9

3 1.60F 0.0110 2.00F 0.776 2.12F 0.0605 0.0402F 0.0003 0.208 8.1 38.8 99.4 959.3F 6.6 Mean Age:

977F 12 Ma

4 1.64F 0.0136 � 1.07F 30.47 2.89F 0.204 0.0429F 0.0003 0.199 6 30.4 100.3 948.2F 7.8

5 1.58F 0.0137 0.35F 0.367 2.38F 0.0588 0.0410F 0.0003 0.201 7.1 35.2 99.9 973.2F 8.5 Regression:

987F 6

6 1.53F 0.0117 1.54F 0.775 2.20F 0.0508 0.0403F 0.0002 0.198 7.4 37.5 99.5 995.4F 7.6 Initial 40/36:

� 36F 28

7 1.55F 0.0123 2.82F 1.22 2.25F 0.0715 0.0407F 0.0003 0.198 7.3 36.9 99.2 982F 7.8

8 1.56F 0.0176 � 0.82F 7.44 2.97F 0.138 0.0413F 0.0003 0.196 5.6 28.4 100.2 987.1F11.2

9 1.63F 0.0172 1.91F1.67 2.54F 0.0970 0.0400F 0.0005 0.212 6.8 32.2 99.4 946.2F 10

10 1.56F 0.0158 1.65F 0.952 2.32F 0.0561 0.0407F 0.0003 0.2 7.2 36 99.5 978.4F 9.9

Mean age calculated with all data (n= 10), as the mean of air-corrected data, with error expressed as the standard error of the mean.

‘‘Regression’’ result is based on the methods of York (1969).

All uncertainties are 2r. The uncertainty in the statistical ages include uncertainties arising from the J-value (0.01132F 0.00006), which

corresponds to an additional 0.5%.

%40Ar*: percentage of radiogenic 40Ar from total 40Ar.a Ages are calculated for each analysis on the basis of analytical precision only.

Notes to Table 1:

abr = abraded; grn = grains; rnd = round; sng = single; prsm= prismatic; cir = circular; pnk = pink; fr = fragment; crk =with cracks; irg = irregular

elg = elongate; sph = spherical; multif =multifaceted; brk xls = broken crystals; CL= imaged by cathodoluminescence; plkd = plucked off from

the CL mount; mg =multigrain; eq = equant.a Denotes radiogenic Pb. Zircon sample dissolution and ion exchange chemistry modified after Krogh (1973) and Mattinson (1987) in

Parrish (1987)-type microcapsules.b Observed isotopic ratios are corrected for mass fractionation of 1x for both 208Pb and 205Pb spiked fractions. Fractions spiked with the

mixed 235U/205Pb tracer are also corrected for spike and blank relative contributions. Uncertainties in the 206Pb/204Pb ratio vary from 0.1% to

2.4%.c Decay constants used: 238U=1.55125�10�10; 235U=9.48485�10�10; 238U/235U=137.88. Estimated uncertainties of the U/Pb ratio are

F0.4 based on replicate analyses of a single zircon standard fraction (see Lopez et al., 2001).d 207Pb*/206Pb* age uncertainties are 2r and from the data reduction program PBDAT of K. Ludwig (1991). Total processing Pb blank

amount varied between 2 and 30 pg, generally averaging <10 pg. Initial Pb composition are from isotopic analysis of feldspar separates or from

the two-stage Pb evolution curve of Stacey and Kramers (1975). Isotopic data were measured on a VG 54-30 sector multicollector mass

spectrometer with a pulse counting Daly detector at UC Santa Cruz.


4.1. Analysis of Zapotecan structures in the Huitzo

unit (magmatic bands, S0, and Z1h)

4.1.1. S0 and Z1hThe anorthosite of the Huitzo unit is weakly

foliated-massive in the middle of the body where a

banding defined by thin ilmenite and apatite bands

may be observed, and was probably produced by

liquid immiscibility and is therefore magmatic in

origin. This S0 banding is deformed by close to

isoclinal, FZ1h intrafolial folds that plunge gently to

moderately towards NW (Fig. 2). They are generally

Fig. 2. Stereoplots obtained for foliations, mineral and stretching lineations and fold axes in the studied area, divided for each structural unit

discussed in the text.


less than 20 cm in amplitude and plot in the Class 3

field, close to Class 2 (Fig. 3a), which suggests they

formed by a buckling mechanism, possibly a combi-

nation of flexural and tangential– longitudinal strain

that was accompanied by flattening. A poorly devel-

oped axial plane foliation defined by rare ortho- and

clinopyroxenes, as well as hornblende is parallel to

the FZ1h axial planes in the FZ1h hinges. In the

margins of the body, the S0 banding is transposed

into the SZ1h foliation represented by dimensionally

Fig. 3. (a–d) tVvs. a diagrams for studied folds associated with different phases of deformation, according to Ramsay (1967). tVis defined as the

ratio between the value ta (thickness between the tangents to two folded surfaces at an angle a p 0), and t0 (thickness between the tangents to

two folded surfaces at the angle a= 0, generally in the hinge zone). a is the angle of dip of the tangents. See Ramsay (1967, p. 358–372), for

further explanations on the dip-isogons classification of folds.


oriented orthopyroxene, clinopyroxene (augite),

brown hornblendeF titaniferous biotite, in the mafic

gneiss and metagabbro. Orthopyroxene constitutes a

relict phase, surrounded by brown hornblende coro-

nas. The gently to moderately NW plunging LZ1h

lineation is defined by oriented hornblende that is

generally parallel to the long axes of quartz ribbons,

which have aspect ratios of up to 15:1. This mineral

association orthopyroxene–clinopyroxene–brown

hornblende indicates that these fabrics formed under

granulite facies conditions. Some of the gabbros show

retrogression at upper amphibolite conditions. Alman-

dine garnet, when present, shows poikilitic textures

and is in textural equilibrium with both pyroxenes and

hornblende. SZ1h poles are distributed along a great

circle whose gently NNW-plunging pole is approxi-

mately parallel to both the FZ1h fold axes, the NW

plunging LZ1h lineations and the Phanerozoic fold axes

(Fig. 2).

4.1.2. Age constraints on the Zapotecan structures in

the Huitzo unit

Ten 40Ar/39Ar laser fusion analyses were per-

formed on hornblende separated from a 100-m-thick

gabbro (sample OC9901) interleaved with the anor-

thosites within the Huitzo unit (Fig. 1b,c). The horn-

blende is in stable contact with augite, plagioclase and

biotite. The analyses are characterized by very high

radiogenic yields (about 99.5% or higher) and there is

no evidence of extraneous, nonatmospheric argon.

Regression of all the data yields an age of 987F 6

Ma, whereas air-corrected data give a mean age of

977F 12 Ma (Table 2). The latter is viewed as

preferable because there is no clear evidence of an

extraneous argon component. This age is interpreted

to date cooling through the argon blocking of f 500

jC temperature in hornblende (Harrison, 1981), and

provides a younger limit on the granulite facies meta-

morphism.

4.2. Analysis of Olmecan (h1c and h2c) and Zapo-

tecan (Z1c and Z2c) structures in the El Catrın unit

4.2.1. h1cThe oldest structures in the El Catrın unit were

only discerned in a strain window (about 500-m

wide) located at the type locality along the Federal

Road 190 (Fig. 1b). Here, migmatitic structures are

preserved, which are progressively transformed in

granulite facies, striped gneisses toward the mar-

gins. Subsequent intense reequilibration under green-

schist facies conditions affected the entire migmatitic

unit. The first recognizable structure in the El Catrın

unit is represented by the stromatic to nebulitic,

centimetric banding of mesosome and leucosome,

Su1c (Fig. 4a,b). We interpret the thin Su1c banding

as the product of metamorphic differentiation and

migmatization that overprinted an original, banded

magmatic feature. The gently NNW plunging nor-

mal to the great circle distribution of the Su1cfoliation poles in the stereonet of Fig. 2 is coinci-

dent with the axes of Fu2c folds and Phanerozoic

fold axes.

4.2.2. h2cThe Su1c banding in El Catrın unit is deformed by

isoclinal, ghostly looking, gently–moderately N-dip-

ping folds (Fu2c) (Fig. 4a,b). These < 15-cm-ampli-

tude folds have gently N plunging hinges that are

parallel to a lineation (stereonet in Fig. 2), which is

constituted by moderately NNW-plunging stretched

quartz ribbons with elongation of up to 8:1. This

lineation probably reflects composite finite strain

accumulated during the Olmecan deformation and

the Z1c event.

4.2.3. Z1cFZ1c folds deform both Su1c migmatitic banding

and Fu2c folds (Fig. 4a). FZ1c folds are up to 20 cm in

size, open to close, gently to moderately N–NNE

plunging, with moderately to gently NE-dipping axial

planes (Fig. 2). Clinopyroxene and orthopyroxene

crystals in the limbs of these folds are aligned

parallel to the FZ1c axial plane, indicating that they

formed under granulite facies conditions but no axial

plane foliation was observed in the hinges of these

folds.

4.2.4. Z2cA SZ2c foliation, which is oblique to the migma-

titic banding (Su1c) and to the FZ1c axial trace, is

microscopically defined by green hornblende that

coexists with quartz and plagioclase (Fig. 4b). It is

parallel to the axial planes of V 40-cm amplitude,

gently W-plunging, gently NW-dipping, open to iso-

clinal, FZ2c folds (Fig. 4b and stereonet of Fig. 2).


The presence of axial planar green hornblende, not

associated with pyroxenes, suggests that FZ2c folds

formed under amphibolite facies metamorphic con-

ditions.

4.2.5. Age constraints on structures in the El Catrın

unit

A total of 16 analyses were performed by U–Pb

geochronology, seven on single or few zircon grains

Fig. 4. (a) Picture illustrating the classic aspect of El Catrın migmatite in the type locality. Two phases of high-grade folding, marked by long

(Fu2c) and short (FZ1c) white arrows, affect the Su1c migmatitic banding. (b) Z-shaped, close to recumbent, amphibolite facies FZ2c fold in El

Catrın migmatite. White lines underline the SZ2c foliation in the outcrop.


separated from the paleosome (67A98), and nine from

the neosome (67B98) (Table 1). All but one yielded

slightly discordant results, a reflection of the internal

complexity of these zircons (Fig. 5, images a, b and

c). The only concordant point yielded an age of

1106F 6 Ma (analysis 7 in Table 1), and was obtained

from a single, abraded tip broken off from a prismatic,

euhedral zircon from the neosome. Its internal cath-

odoluminescence image (image d in Fig. 5) shows a

thin, oscillatory zoning, which suggests magmatic

growth. We interpret this concordant age of 1106F 6

Ma as the best estimate of the time of migmatization. A

high luminescent overgrowth (possibly metamorphic

in origin) was removed by abrasion, performed after

the zircon was plucked off from the CL mount. Several

chords may be drawn: (i) through all the data, yielding

intercepts of 1358F 43 and 1048F 31 Ma, with an

MSWD of 10.6; (ii) through the concordant point at

1106 Ma and the older data, yielding intercepts of

1399F 58 and 1092F 38 Ma, with an MSWD of 7.3

(shown in Fig. 6a); and (iii) through the concordant

point at 1106 Ma and the younger data, yielding inter-

cepts at 1116 + 50/� 27 and 937F 100 Ma. The upper

intercepts in (i) and (ii) of 1358F 43 and 1399F 58

Ma, respectively, are inferred to provide a minimum

age for the protolith. On the other hand, the lower

intercepts in (i) and (iii) of 1048F 31 and 937F 100

Ma, respectively, may provide a rough estimate of the

age of the granulite metamorphism in the El Catrın

unit.

In the upper part of the El Catrın unit, a pegmatite

(6898) folded by a FZ1c fold (Fig. 1b) is composed of

quartz, plagioclase (An 20%) and alkalic feldspar, with

accessory zircon, apatite and secondary epidote and

chlorite. Zircons separated from this pegmatite are

elongated, clear, light pink coloured. Cathodolumines-

cence images reveal internal homogeneity and thin,

low luminescent overgrowths in the elongate zircons

(image e in Fig. 5), and thin, oscillatory zoning in the

stubby to elongate zircons, which suggests magmatic

crystallization. A chord through two slightly discordant

analyses (Table 1) yields an upper intercept at 1126 +

11/� 6Ma (Fig. 6b), which is tentatively interpreted as

a minimum crystallization age for this pegmatite.

Fig. 5. Cathodoluminescence images for some of the dated zircons. The white bar close to each image represents 100-Am scale. (a–d) Zircons

from the El Catrın migmatite 6798; (e–g) zircons from the pegmatite 6898; (h) zircon from the pegmatite 66D98.


Fig. 6. U–Pb concordia plots. (a) El Catrın migmatite, 67A98 and 67B98; (b) El Cuajilote folded pegmatite, 6898; (c) semipelitic metasediment,

7098; (d) pre-tectonic pegmatite, 66B98; (e) flattened, axial planar pegmatite, 66C98; (f) Late tectonic pegmatite, 66D98; (g) El Tecolote

pegmatite, 6998. Errors are quoted at F 2r.


4.3. Analysis of Olmecan (h1m and h2m) and

Zapotecan (Z1m and Z2m) structures in the El

Marquez unit

4.3.1. h1mThe V 5 mm Su1m banding consists of alternating

leucocratic and mafic horizons that could have formed

by metamorphic differentiation that may, or may not,

have enhanced an original fabric in the rocks. In the

orthogneisses, the leucocratic bands are made up of

various proportions of quartz, plagioclase, potassic

feldspar and, sometimes, almandine garnet. The mafic

horizons are composed of orthopyroxene, clinopyrox-

ene, hornblende, biotite, plagioclase and opaque min-

erals. In the paragneisses, the leucocratic bands also

contain sillimanite, rutile and graphite. The foliation,

Su1m, is defined by planar distribution of aligned

hornblende and/or titaniferous biotite, and stretched

quartz, whereas in the meta-syenite and charnockite, it

is defined by aligned orthopyroxene, clinopyroxene

and hornblende aggregates, which are parallel to

stretched quartz ribbons. Only one mineral lineation

was observed in the foliation and because the rocks

have been deformed several times at high grade

during the Olmecan and Zapotecan events, it is

inferred that the lineation is the finite result of several

increments of strain. This composite, NW to N

plunging Lu1–2m–Z1m mineral lineation in the orthog-

neisses is defined by the long axes of pyroxene, mica

and hornblende, and in the paragneisses by mica and

orthopyroxene is parallel to a stretched quartz line-

ation. The Su1m foliation poles are generally distrib-

uted along a great circle, whose axis is parallel to the

Fu2m isoclinal folds axes, the FZ1m fold axes and

Phanerozoic fold axes (Fig. 2).

4.3.2. h2mThe Su1m foliation and banding is deformed by

close to isoclinal, gently to moderately NNW to N

plunging, gently NNW dipping, S and Z asymmetric

folds (Fig. 2). The folds are common in paragneisses,

and only a few were observed in the orthogneisses.

Rootless fold hinges in quartz-feldspar veins were

also observed in mica-rich bands. The Fu2m folds are

Class 1B–1C in the charnockites, 1C in mafic

horizons and 1C close to 2 in metasediments (exam-

ples in Fig. 3b). These data may be interpreted in

terms of a combination of flexural and tangential–

longitudinal strain mechanisms, with superposed flat-

tening. There is no a well-developed axial plane

foliation and pyroxene and hornblende are stable in

the fold hinges.

4.3.3. Z1mGently NNW plunging, gently N dipping, FZ1m

close folds up to V 10 m are common in paragneisses

(stereonet in Fig. 2). They range from Class 1C to 3,

close to Class 2 (Fig. 3c). This can be interpreted as a

combination of flexural and tangential–longitudinal

strain mechanism, affected by flattening perpendicular

to the axial plane. At one location, a gently folded

pegmatite cuts across both limbs of a nearly isoclinal

fold and allows the strain to be partitioned into pre-

and post-intrusion components (Fig. 7). The pegma-

tite contains flattened discs of quartz, whereas the host

gneisses contain high-stretched quartz. The flattening

component can be estimated using the methods pro-

posed by Lisle (1992) or Weger (1993). An oriented

sample of this pegmatite was cut perpendicular to the

fold axis, polished to expose the SZ1m surface, then

photographed and finally restored to allow the flat-

tened quartz to recover a quasi-equidimensional shape.

Both methods give a two-dimensional ratio of 1:3.3,

which corresponds to the strain component due to

flattening superimposed on FZ1m folding. Quartz–

plagioclase – garnetF orthopyroxene association

observed along the axial plane of one of these folds,

as well as the paragenesis garnet–sillimanite–rutile–

quartz, suggest FZ1m folds formed under granulite

facies metamorphism.

4.3.4. Z2mThese structures consist of gently N-plunging,

steeply inclined FZ2m folds (Fig. 2). They show a

well-developed axial plane foliation in the hinge zone

defined by hornblende crystals that crosscut the Su1mbanding (Fig. 8a). The outer arc of one of these folds

is intruded along its axial plane by a pegmatite

(sample 66C98) that displays flattened quartz in the

margins, whereas its centre is not foliated (Fig. 9b).

The shape of the outer arc of the fold changes from

Class 1A to 1C with the increasing of curvature,

whereas the inner arc of this fold is Class 3 close to

2 (Fig. 3d). These data suggest a tangential– longitu-

dinal strain mechanism, where the outer arc of the fold

hinge was stretched parallel to the banding facilitating


Fig. 7. Granulite facies, mesoscale recumbent FZ1m fold, which limbs are cut across by a 15-cm-thick pegmatite, unfolded but affected by

flattening perpendicular to the fold axis. The estimated flattening corresponds to an S1/S2 ratio of f 3.3:1. White lines in the detail (b) marks

the axial plane foliation in both pegmatite and folded host rock. Scale in picture (b) is 5-cm large. See text for further explanations.


Fig. 8. (a) Photomicrograph taken in an oriented thin section, showing the SZ2m hornblende retrogression foliation, cutting across the previous

hornblende (Hbl1, basal sections) + augite (Cpx) + plagioclase (Pl) banding of the hosting folded metagabbros. The dark, new hornblende (Hbl2)

grows in the SZ2m planes, roughly oriented NW–SE in the picture. The black arrows indicate the triple joints in the plagioclase matrix. Parallel

polars. The long side of the picture corresponds to 2.5 mm. (b) Photomicrograph of the flattened quartz in the margins of the pegmatite 66C98,

which shows microstructures such as subgrains with domino-like shape and 90j angles, with undulose extinction and internal, oblique

deformation lamellae (white arrows). The flattened quartz grains make an angle of f 10j with respect of the pegmatite margins. Crossed

polars. The long side of the picture corresponds to a 2.5-mm field.


intrusion of the pegmatite. It is inferred that continued

deformation led to flattening of the quartz in the

margins of the pegmatite, which shows ductile micro-

structures such as (a) formation of subgrains with

domino-like shape and 90j angles, (b) undulose

extinction and (c) internal, oblique deformation lamel-

lae (Fig. 8b). The hornblende parallel to the axial

plane foliation and adjacent to the pegmatite suggest

that the fold formed under amphibolite facies con-

ditions, and that fluids necessary to hydrate the

Fig. 9. (a) Picture of the outcrop at km 199.2 of the Highway Mexico City-Oaxaca, where the samples 66A98, 66B98, 66C98 and 66D98 were

collected. Each sample, together with the U–Pb interpreted age, is indicated. Sample 66A98 is dated and discussed in Keppie et al. (2003). (b)

Schematic of the 66C98 pegmatite field relationships with respect to the hosting FZ2m fold. See text for further explanations.


clinopyroxene, stable far from the pegmatite intrusive

contacts, were probably provided by the pegmatite.

4.3.5. Age constraints on structures in the El Marquez

unit

A sample of metasediment in the El Marquez-type

locality (sample 7098) was analysed to constrain its

protolith and granulite metamorphism age. The

selected sample has a semipelitic composition, being

made up of quartz, plagioclase, graphite, biotite, relict

pyroxene, opaque minerals, and accessory zircon and

apatite. Three analyses were performed on selected

grains. Analysis 19 on an equant and abraded single

crystal yielded a concordant age of 994F 3 Ma (Table

1), which is interpreted as recording the granulite

facies metamorphism. Analyses 20 and 21 yielded

discordant data (Table 1). A chord drawn through all

three analyses in the concordia diagram of Fig. 6c

yields a lower intercept of 996 + 18/� 22 Ma and an

upper intercept of 1300 + 69/� 61 Ma, which is

possibly a mixing of detrital zircons and the f 1106

Ma Olmecan event.

In order to bracket the structural fabrics in the El

Marquez unit, several samples (66B98, 66C98,

66D98) were collected from early-, middle- and late-

stage pegmatites cutting a metagabbro on the Federal

Highway at 199.2 km (Fig. 9a). Zircons in the

metagabbro yielded discordant analyses that fall on

a chord between 1257F 71 and 1021F 39 Ma, which

are interpreted as the minimum protolith age and as a

rough estimate of the age of granulite facies meta-

morphism, respectively (Keppie et al., 2003). Sample

66B98 is a 2-m-thick, foliated pegmatite that is nearly

concordant with the foliation in the surrounding

metagabbro. It is composed of quartz, alkalic feldspar,

plagioclase, hornblende, magnetite and apatite. U–Pb

analyses yielded one concordant result with an age of

1144F 2 Ma on an abraded three-grain fraction inter-

preted as a minimum for the time of intrusion (Table

1, analysis 23), and several � 0.4% to + 1.7% dis-

cordant points. All of these analyses fall on a chord

with a MSWD of 5.9, with upper and lower intercepts

of 1144F 36 and 925F 200 Ma; the latter is within

error of the time of granulite facies metamorphism.

The 1144F 2 Ma age is similar to the 1131F10 Ma

upper intercept U/Pb age determined for the meta-

syenite in the El Marquez unit (Keppie et al., 2003). It

is also within error of the minimum crystallization age

of 1126 + 11/� 6 Ma calculated on the pegmatite

6898 in the El Catrın structural unit, suggesting that

they may be related intrusions.

The pegmatite 66C98 intruding the outer arc of the

FZ2m fold described above (Fig. 9b) is made up of

quartz, pink perthitic feldspar, plagioclase (An 25%),

rare hornblende, magnetite and ilmenite. Zircons from

this pegmatite vary from pink to red, stubby to broken,

clear, with rounded tips and no visible cores. Five

analyses were performed on unabraded/abraded multi-

grain fractions and yielded concordant to slightly

discordant ages with 207Pb/206Pb between 971 and

984 Ma (Table 1). Analysis 30 was performed on

seven unabraded stubby grains with sharp edges

between the facets, and yielded a concordant age of

978F 3 Ma, which is inferred to represent the time

of intrusion. Because this pegmatite intrudes the

stretched, outer arc of a FZ2m fold and is itself ductilely

deformed during the last increment of flattening, we

interpret the concordant age of 978F 3 Ma as provid-

ing the best estimate of the age of the FZ2m folding

event formed under amphibolite facies conditions.

An undeformed, massive, 50-cm-thick pegmatite

(66D98) that cuts across the foliation in the metagab-

bro host rock and the foliated 66B98 pegmatite (Fig.

9a) was sampled for U/Pb geochronology. It consists

of quartz, pink perthitic feldspar, microcline, plagio-

clase (An 25%), magnetite, retrogressed hornblende,

and apatite and zircon as accessory minerals. Four,

abraded, single zircons were analysed for U–Pb

isotopes and yielded concordant to slightly discordant

data (Table 1), with the weighted mean giving an age

of 977F 2 Ma (Fig. 6f). Analyses 33 and 34 in Table

1 were performed on stubby to rounded grains. A

similar, stubby zircon was imaged by cathodolumi-

nescence (image h in Fig. 5) and its internal morphol-

ogy is characterized by thin concentric oscillatory

zoning, suggesting a magmatic origin.

A 2-m-thick pegmatite (6998) intrudes quartz–

feldspar–garnet paragneisses and cuts across the

Su1m foliation. Thin pegmatitic dikes emanating from

the main body contain a C–S fabric of stretched

quartz with top to SE sense of shearing, which are

inferred to have formed at lower amphibolite–upper

greenschist facies metamorphic conditions (Passchier

and Trouw, 1996). This pegmatite is composed of

bluish quartz, plagioclase, perthitic feldspar and very

large zircons, up to 2 cm in size. The intense blue


Fig. 10. Summary of the tectonic events recognized in the three tectonic units of the northern Oaxacan Complex, and discussed in the text. The intrusive events reported by Keppie et

al. (2003) are shown in italic.

L.A.Solariet

al./Tecto

nophysics

365(2003)257–282

275

colour in the quartz is probably due to microscopic

inclusions of Ti–Fe oxides, such as rutile or ilmenite.

This scattering phenomenon has been described else-

where in high-grade gneiss terrains (e.g., Herz and

Force, 1984), and it is inferred to be due to a high-

grade ilmenite exolution. Thus, although the pegmatite

was probably intruded under granulite facies condi-

tion, it was deformed at a much lower grade of meta-

morphism. Zircons occur as big, dark purple euhedral

prismatic crystals. One large zircon was broken into

fragments, which were intensely abraded, and then two

of them were chosen for U–Pb geochronology. The

first yielded a 0.5% discordant data with 207Pb/206Pb

age of 975F 1 Ma (Table 1, analysis 37). The second

fragment yielded an almost concordant age (� 0.1%

reverse discordance) at 979F 3 Ma (Table 1, analysis

38), which is interpreted as the time of pegmatite

crystallization. If the blue quartz formed under gran-

ulite facies conditions, then the 979F 3 Ma age

records the last stage of this grade of metamorphism.

4.4. Summary

Using these data together with that in Keppie et al.

(2003), all the Zapotecan granulite and upper amphib-

olite fabrics are bracketed between the 1012F 12 Ma

protolith age in the Huitzo unit and the 977F 2 Ma

age of the post-tectonic pegmatite (66D98). Granulite

facies metamorphic ages range between the concord-

ant 1004F 3 Ma age in charnockite of the Huitzo unit

(Keppie et al., 2003) and the concordant 979F 3 Ma

age on the post-tectonic, blue quartz pegmatite

(6998). Using all the data, the transition from gran-

ulite to amphibolite facies conditions took place

between 981 and 976 Ma: the overlap between

979F 3 Ma (sample 6998: pegmatite intruded during

granulite facies metamorphism) and 978F 3 Ma

(sample 66C98: pegmatite intruded during upper

amphibolite facies of metamorphism), respectively.

Following the Zapotecan tectonothermal event, cool-

ing through f 500 jC occurred quickly, as indicated

by the 977F 12 Ma age on the hornblende.

4.5. Correlation of structures in the three units

Intrusion of the AMCG units at f 1012 Ma of the

Huitzo unit clearly separates the f 1100 Ma Olme-

can and f 1004–980 Ma Zapotecan tectonothermal

events, and makes identification of the Zapotecan

structures relatively clear. Thus, we correlate the

Zapotecan structures in the Huitzo unit (Z1h: NW-

plunging lineation, a generally NW-dipping foliation,

and NW-plunging, northerly inclined folds) with sim-

ilarly oriented structures in the El Catrın (Z1c) and El

Marquez units (Z1m) (Figs. 2 and 10). This is con-

sistent with their formation under granulite facies

metamorphic conditions. The structures formed under

upper amphibolite facies metamorphic conditions in

the El Catrın (Z2c) and El Marquez (Z2m) units have

distinctly different orientations from the Z1 structures:

W-plunging nearly recumbent folds in the El Catrın

unit, and NNE-plunging, upright folds in the El

Marquez unit (Figs. 2 and 10).

This leaves two earlier sets of structures in the El

Catrın and El Marquez units, which are assigned to

the Olmecan tectonothermal event. Thus, the mig-

matization in the El Catrın unit (u1c) may be broadly

synchronous with metamorphic differentiation in the

El Marquez unit that produced the banding (u1m)

(Figs. 2 and 10). This allows the similarly oriented

isoclinal folds (u2) in these two units to be corre-

lated (Figs. 2 and 10). The metamorphic assemb-

lages formed during the Olmecan event were

overprinted by the Zapotecan granulite facies meta-

morphism.

5. Discussion

Data presented here indicate that two main tecto-

nothermal events occurred in the northern Oaxacan

Complex of southern Mexico: 1106F 6 Ma Olmecan

event and f 1004–978F 3 Ma Zapotecan event.

The significance of the Olmecan migmatization and

deformation is uncertain, and it has yet to be recorded

elsewhere in the f 1.1 Ga rocks of Mexico. The

Olmecan event is approximately contemporaneous

with a rift-related granite dated at 1117F 4 Ma in

the southern Oaxacan Complex, and follows rift-

related magmatism at f 1140 Ma (Keppie et al.,

2001, 2003). Thus, it is possible that the Olmecan

event is related to rifting, in which case it may be

extensional rather than compressional. Assuming a rift

model, partial melting could have been triggered by a

rise in the isotherms caused by rifting. Clearly, more

data are required, such as kinematic data and more


conventional U–Pb and SHRIMP dating, to better

define the nature and significance of the Olmecan

event.

The Zapotecan tectonothermal event in the north-

ern Oaxacan Complex is recorded by widespread

polyphase deformation and associated granulite facies

metamorphism. A contemporaneous tectonothermal

event has also been recorded in all the other Mexican

inliers that form Oaxaquia, such as in the Guichicovi

Complex (f 986 Ma, Ruız et al., 1999; Weber and

Kohler, 1999), in the Huiznopala Gneiss (f 1000

Ma, Lawlor et al., 1999), and in the Novillo Gneiss

(f 928–1018 Ma, Denison et al., 1971; Silver et al.,

1994). Although a contemporaneous event has not

been recorded in Texas (Mosher, 1998 and references

therein) or in most of the Grenville Province, the

Zapotecan event is contemporaneous with the 980–

1000 Ma Rigolet event, which is restricted to amphib-

olite–greenschist facies reactivation along the Gren-

ville Front (Scharer et al., 1986; Krogh, 1994;

Corrigan et al., 2000). Several Grenvillian massifs in

the southern Appalachians also record a f 990–1010

Ma tectonothermal event (references in Sinha et al.,

1996). Aleinikoff et al. (1996) reported a granulite

Fig. 11. 1 Ga Rodinia reconstruction (modified from Keppie et al., 2003) showing 990–1000 Ma tectonothermal events in the Grenville-aged

orogens (dark, diagonal pattern). Bold numbers 1 to 3 indicate the possible locations of Oaxaquia (OX) and Chortis Block (CH), as discussed by

Keppie et al. (2003). Abbreviations: SM=Santa Marta Massif; AA=Arequipa–Antofalla Massif; R =Rockall Plateau (Goochland and Carolina

terranes); SN = Sveco–Norwegian Orogen; CM=Coats Land/Maudheim/Grunehogna terrane; SO= Sunsas Orogen; WA–BR=West

Avalonia–Blair River; SF = Sao Francisco Craton.


Fig. 12. Schematic cartoon illustrating the evolution of Oaxaquia during the Zapotecan orogeny (f 990 Ma). (a) Andean-type orogen, in which

retro-arc thrusting places the Avalonian magmatic arc over the Oaxaquia backarc where rift-related AMCG magmatism occurred at f 1012 Ma:

the Andes would be the modern analogue; (b) arc–continent collision model, where the overriding Avalonian arc collides with Oaxaquia, the

passive margin of the Amazonian craton: Australia–Papua New Guinea would be the modern analogue; (c) continent–continent collision

model, in which a continent (Laurentia?) overrides the passive margin of Amazonia constituted by Oaxaquia: the Alps can be viewed as modern

analogue.


facies metamorphism at f 1011 Ma of the f 1045

Ma Montpellier anorthosite in the Goochland terrane

of the southern Appalachians. Similarly, the upper

amphibolite facies metamorphism at f 996 Ma in the

Blair River Complex of the northern Appalachians is

synchronous with the Zapotecan event (Miller et al.,

1996; Miller and Barr, 2000), as well as the 980–960

Ma Sveconorwegian Orogeny of southwestern Baltica

(Soderlund et al., 2002). Restrepo-Pace (personal

communication, 2000) reported a tectonothermal

event at 981F 85 Ma in the Colombian massifs of

the Andes. Wasteneys et al. (1995) reported a 970F 23

Ma, granulite facies metamorphism and deformation in

the Mollendo Domain of the Arequipa Massif of Peru.

Although Litherland et al. (1986) reported a f 990

Ma tectonothermal event in the Sunsas Orogen of

Bolivia, recent U–Pb data suggests that it is consid-

erably older (Tassinari et al., 2000). Thus, the distri-

bution of granulite facies metamorphism during the

f 1 Ga tectonothermal event appears to be restricted

to the central parts of the Grenville Orogen in a Rodi-

nia reconstruction (Fig. 11).

The Zapotecan granulite facies tectonothermal

event that occurred at depths of f 30 km probably

formed in the roots of an orogen produced during

subduction and/or collision (Jamieson et al., 1998).

We consider three models (Fig. 12): (1) fold-and-

thrust belt associated either with A-subduction (e.g.

the Andes: Ramos and Aleman, 2000), or with shal-

lowing of the subduction zone, subduction of young

oceanic lithosphere, and overriding of plumes (e.g.

Laramide Orogeny: Murphy et al., 1998); (2) arc–

continent collision (e.g. Papua/New Guinea and Aus-

tralia: van Staal et al., 1998); and (3) continent–

continent collision (e.g. Alps: Stampfli et al., 1998).

Although polyphase deformation is present in all these

tectonic settings, a contemporaneous association with

granulite facies metamorphism is perhaps more diag-

nostic. England and Richardson (1977) and Jamieson

et al. (1998) have shown that >25–40 million years is

required to reequilibrate the isotherms following crus-

tal thickening. This problem may be overcome by

delamination of the orogenic root, which is replaced

by upwelling hot asthenosphere associated with mag-

matism; however, the latter is not present in the

Zapotecan. Delamination has been proposed in sev-

eral tectonic settings including the Andes, arc–con-

tinent and continent–continent collisions and so

cannot be used to discriminate between them. Syn-

chroneity of granulite facies metamorphism and defor-

mation may also be overcome by overriding a plume

as occurred in the Laramide orogeny; however, here;

it is associated with widespread magmatism (Murphy

et al., 1998), a feature apparently absent in Oaxaquia

during the Zapotecan. On the other hand, is it possible

that intrusion of the f 1012 Ma AMCG suite pro-

vided a sufficiently large initial temperature increment

that survived into the Zapotecan deformation episode?

Even with a positive response to this question, one

cannot discriminate between the various tectonic set-

tings.

Thus, all three tectonic settings envisaged for the

Zapotecan are possible: (1) continent–continent colli-

sion between two land masses that could be Laurentia

and Amazonia (Fig. 12a); (2) arc–continent collision

between the Avalonia–Carolinia juvenile arc and

Oaxaquia (Fig. 12b); or (3) an Andean-type orogen

adjacent to the northern margin of Amazon craton

(Fig. 12c).

Acknowledgements

We gratefully acknowledge funding for various

aspects of this project from CONACyT grants (0225P-

T9506 and 25705-T), PAPIIT-UNAM grants

(IN116999 and IN107999) to JDK and FOG, a UC-

MEXUS grant to KLC and FOG, NSF grant

EAR9909459 to KLC and UNAM-PAEP student grant

to LAS. We would also like to thank Pete Holden for

assistance with isotopic analyses at UCSC, and Kip

Hodges for access to the MIT 40Ar/39Ar facility. Care-

ful reviews of I. Fitzsimons, E. Johnson and P. Schaaf

greatly improved the clarity of the manuscript, and

helped to clarify some points. This work is a

contribution to the IGCP # 453 (Modern and Ancient

Orogens).

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