The North American–Caribbean Plate boundary in Mexico ... · ocean-basin deposits (North El Tambor complex and El Pilar Formation, Fig. 2a) and crystalline basement of the Chuacu´s

The North American–Caribbean Plate boundary in

Mexico–Guatemala–Honduras

LOTHAR RATSCHBACHER1*, LEANDER FRANZ1, MYO MIN1, RAIK BACHMANN1,

UWE MARTENS2, KLAUS STANEK1, KONSTANZE STUBNER1, BRUCE K. NELSON3,

UWE HERRMANN3, BODO WEBER4, MARGARITA LOPEZ-MARTINEZ4,

RAYMOND JONCKHEERE1, BLANKA SPERNER1, MARION TICHOMIROWA1,

MICHAEL O. MCWILLIAMS2, MARK GORDON5, MARTIN MESCHEDE6

& PETER BOCK1

1Geowissenschaften, Technische Universitat Bergakademie Freiberg, 09599 Freiberg, Germany2Geological and Environmental Sciences, Stanford University, Stanford, CA 94305, USA

3Earth and Space Sciences, University of Washington, Seattle, WA 98195, USA4CICESE, 22860 Ensenada, B.C., Mexico

5Department of Geology and Geophysics, Rice University, Houston, TX 77251-1892, USA6Geowissenschaften, Universitat Greifswald, 17487 Greifswald, Germany

*Corresponding author (e-mail: [email protected])

Abstract: New structural, geochronological, and petrological data highlight which crustal sec-tions of the North American–Caribbean Plate boundary in Guatemala and Honduras accommo-dated the large-scale sinistral offset. We develop the chronological and kinematic frameworkfor these interactions and test for Palaeozoic to Recent geological correlations among the MayaBlock, the Chortıs Block, and the terranes of southern Mexico and the northern Caribbean. Ourprincipal findings relate to how the North American–Caribbean Plate boundary partitioned defor-mation; whereas the southern Maya Block and the southern Chortıs Block record the Late Cretac-eous–Early Cenozoic collision and eastward sinistral translation of the Greater Antilles arc, thenorthern Chortıs Block preserves evidence for northward stepping of the plate boundary withthe translation of this block to its present position since the Late Eocene. Collision and translationare recorded in the ophiolite and subduction–accretion complex (North El Tambor complex), thecontinental margin (Rabinal and Chuacus complexes), and the Laramide foreland fold–thrust beltof the Maya Block as well as the overriding Greater Antilles arc complex. The Las Ovejas complexof the northern Chortıs Block contains a significant part of the history of the eastward migration ofthe Chortıs Block; it constitutes the southern part of the arc that facilitated the breakaway of theChortıs Block from the Xolapa complex of southern Mexico. While the Late Cretaceous collisionis spectacularly sinistral transpressional, the Eocene–Recent translation of the Chortıs Block is bysinistral wrenching with transtensional and transpressional episodes. Our reconstruction of the LateMesozoic–Cenozoic evolution of the North American–Caribbean Plate boundary identifiedProterozoic to Mesozoic connections among the southern Maya Block, the Chortıs Block, andthe terranes of southern Mexico: (i) in the Early–Middle Palaeozoic, the Acatlan complex ofthe southern Mexican Mixteca terrane, the Rabinal complex of the southern Maya Block, theChuacus complex, and the Chortıs Block were part of the Taconic–Acadian orogen along thenorthern margin of South America; (ii) after final amalgamation of Pangaea, an arc developedalong its western margin, causing magmatism and regional amphibolite–facies metamorphismin southern Mexico, the Maya Block (including Rabinal complex), the Chuacus complex andthe Chortıs Block. The separation of North and South America also rifted the Chortıs Blockfrom southern Mexico. Rifting ultimately resulted in the formation of the Late Jurassic–Early Cre-taceous oceanic crust of the South El Tambor complex; rifting and spreading terminated before theHauterivian (c. 135 Ma). Remnants of the southwestern Mexican Guerrero complex, which alsorifted from southern Mexico, remain in the Chortıs Block (Sanarate complex); these complexesshare Jurassic metamorphism. The South El Tambor subduction–accretion complex was emplacedonto the Chortıs Block probably in the late Early Cretaceous and the Chortıs Block collided with

From: JAMES, K. H., LORENTE, M. A. & PINDELL, J. L. (eds) The Origin and Evolution of the Caribbean Plate.Geological Society, London, Special Publications, 328, 219–293.DOI: 10.1144/SP328.11 0305-8719/09/$15.00 # The Geological Society of London 2009.

southern Mexico. Related arc magmatism and high-T/low-P metamorphism (Taxco–Viejo–Xolapa arc) of the Mixteca terrane spans all of southern Mexico. The Chortıs Block shows continu-ous Early Cretaceous–Recent arc magmatism.

Supplementary material: Analytical methods and data, and sample description are available athttp://www.geolsoc.org.uk/SUP18360.

Structural, geochronological, and petrological datadocument the Palaeozoic–Cenozoic pressure–temperature–deformation–time (P–T–d– t) his-tory of the southern part of the Maya and northernpart of the Chortıs Blocks, and the Polochic,Motagua and Jocotan–Chamelecon fault zonesalong the North American–Caribbean Plate bound-ary in southern Mexico, Guatemala and northernHonduras. The principal questions addressed are:which crustal sections accommodated the large-scale sinistral offset along the plate boundary,what was the timing of offset and deformation,and in what kinematic framework did it occur?What Palaeozoic to Cenozoic geological corre-lations exist between the Maya Block, the ChortısBlock, the tectono-stratigraphic complexes (‘ter-ranes’) of southern Mexico, and the arc/subductioncomplexes of the northern Caribbean.

Northern Caribbean Plate boundary

The Caribbean Plate represents a lithospheric unitbetween the North and South American plates. Itswestern and eastern margins consist of subductionzones with magmatic arcs (Central AmericaIsthmus, Lesser Antilles), whereas the northernand southern margins correspond to strike–slipwith subsidiary transpression or transtension zones(Polochic–Motagua, Cayman, Greater Antilles,North Andean and South Caribbean, e.g. Pindellet al. 2006). The Motagua Suture Zone (MSZ) ofsoutheastern Mexico, Guatemala, and Hondurasextends from the Pacific Ocean to the CaribbeanSea for c. 400 km east–west and c. 80 km north–south (e.g. Beccaluva et al. 1995), separating theMaya Block from the Chortıs Block (Fig. 1a,Dengo 1969). This belt actually represents a compo-site structure, encompassing suture zone(s) withrelics of Mesozoic ocean floor that characterizethe northern Caribbean Plate boundary, and Ceno-zoic fault zones (mainly the Polochic, Motagua,Jocotan–Chamelecon faults) that link the Caymantrough pull-apart basin in the east with the subduc-tion zone of the Pacific plate in the west (Fig. 1).

The Maya and Chortıs Blocks showcase a geo-logical history that encompasses Middle Proterozoicto Quaternary and have long been recognized askey elements in understanding the interactionof Laurentia, Gondwana and the (Proto-)Pacific

domains (e.g. Dickinson & Lawton 2001; Keppie2004). Most tectonic models suggest that collisionof the Maya and Chortıs Blocks occurred alongsouth- or north-dipping subduction zones duringthe Cretaceous, and caused emplacement ofJurassic–Cretaceous Proto-Caribbean oceanic crustonto the blocks; thrusting of metamorphic andsedimentary strata was bivergent north and southof the Motagua fault zone (e.g. Meschede &Frisch 1998; Beccaluva et al. 1995 and referencestherein). Since Late Cretaceous–Early Cenozoic,the MSZ developed into a sinistral wrench zonedue to the different velocities of the North Americanand Caribbean plates (e.g. DeMets 2001). Insouthern Mexico, Guatemala and Honduras thisstill active fault system has dismembered the colli-sional blocks as well as the allochthonous remnantsof the former oceanic crust; the latter comprisesshaly melange, containing blocks of serpentinizedophiolite fragments, and metamorphic rocks suchas amphibolite, eclogite, albitite and jadeitite (e.g.Tsujimori et al. 2005). Inferred Cenozoic displace-ment accommodated along the MSZ is controver-sial, but magnetic anomalies within the Caymantrough suggest a minimum of c. 1100 km; spreadingrates in the trough were c. 3 cm/annum from 44 to30 Ma and c. 1.5 cm/annum thereafter (Rosen-crantz et al. 1988). Active faulting and volcanismrelated to eastward subduction of the Pacific plateoverprinted the ophiolite emplacement and intra-continental translation features within the crust ofMexico, Guatemala and Honduras (e.g. Guzman-Speziale 2001).

Methods

Methodical and technical aspects of our radiometricdating that includes U–Pb, Ar–Ar, Rb–Sr, andfission-track geochronology are provided asSupplementary material. Also included in the Sup-plementary materials is our approach to structuralanalysis of ductile deformation and kinematics,to fault–slip analysis and definition of stress-tensorgroups in the brittle crust, and a review of theapplied calculation techniques. Methodical andtechnical aspects of our petrology and geothermo-barometry are included in the Supplementarymaterial along with locations of studied outcrops(‘stops’) and descriptions of analysed samples.

L. RATSCHBACHER ET AL.220

Fig. 1. (a) Plate tectonic framework of northern Central America and crustal complexes (‘terranes’) of southern Mexico. Numbers in the Cayman trough give age of transitional to oceanic lithosphere (from Rosencrantz et al. 1988). (b) SRTM digital elevation modelof the Motagua suture zone in southern Mexico, Guatemala, and Honduras, highlighting neotectonic structures such as the Polochic, Motagua and Tonala strike–slip faults and NNE-trending grabens (e.g. Sula graben). The Jalpatagua fault zone comprisesdistributed neotectonic dextral strike–slip deformation within the Cenozoic arc. (c) Basement geology of southern Mexico and northern Central America and available reliable geochronological data surrounding the Motagua suture zone. References for Belize,Guatemala and Honduras: 1, Ortega-Gutierrez et al. (2004); 2, Harlow et al. (2004); 3, Horne et al. (1976a); 4, Manton (1996); 5, Venable (1994); 6, Steiner & Walker (1996); 7, recalculated from Gomberg et al. (1968); 8, unpublished, referenced in Donnelly et al.(1990); 9, Bertrand et al. (1978); 10, Manton & Manton (1984); 11, Drobe & Oliver (1998); 12, Donnelly et al. (1990); 13, Emmet (1983); 14, Italian Hydrothermal (1987; reported in Rogers 2003); 15, Horne et al. (1976b); 16, Horne et al. (1974); 17, MMAJ (1980;reported in Rogers 2003); 18, Manton & Manton (1999); 19, Weiland et al. (1992); 20, Bargar (1991); 21, Ortega-Obregon et al. (2008); 22, Solari et al. (2008); 23, Martens et al. (2008). References for Mexico: A, Ortega-Gutierrez et al. (1999); B, Talavera-Mendozaet al. (2005); C, Sanchez-Zavala et al. (2004); D, Yanez et al. (1991); E, Weber et al. (2006a); F, Ducea et al. (2004a); G, Elıas-Herrera & Ortega-Gutierrez (2002); H, Solari et al. (2001); I, Vega-Granillo et al. (2007); J, Keppie et al. (2004); K, Moran-Zenteno(1992); L, Solari et al. (2007); M, Herrmann et al. (1994); N, Elıas-Herrera et al. (2005); O, Tolson (2005); P, Robinson et al. (1989); Q, Wawrzyniec et al. (2005); R, Weber et al. (2005); S, Alaniz-Alvarez et al. (1996); T, Solari et al. (2004); U, Schaaf et al. (2002); V,Weber et al. (2008). Ages without label are our own data from southern Mexico and the central Chortıs Block.

Basement and cover of the Maya and

Chortıs Blocks

Figure 2 presents lithostratigraphic columns of thesouthern Maya Block, the MSZ, and the north-central Chortıs Block, based mainly on the compi-lations given in Donnelly et al. (1990), Talavera-Mendoza et al. (2005, 2007), Vega-Granillo et al.(2007) and Rogers (2003). We modified the columnsaccording to our new data and the recent literature(discussed below) and, where available, added strati-graphic range (converted into Ma), thickness, reli-able radiometric ages, PT estimates and first-ordertectono-stratigraphic interpretations. We also showcolumns of units that we will use in the discussionfor correlation and refinement of tectonic models.

Cover sequences

In general, the Maya Block (Fig. 2b) records shelfevolutions during Late Palaeozoic and Cretaceousthat were followed by terrestrial–shallow marinerift (Middle–Late Jurassic) and submarine-fan(Late Cretaceous) developments. Along the MSZ,ophiolites are imbricated with probably UpperJurassic to Upper Cretaceous submarine-fan andocean-basin deposits (North El Tambor complexand El Pilar Formation, Fig. 2a) and crystallinebasement of the Chuacus complex; they are uncon-formably overlain by continental red beds (SubinalFormation). In the Sierra de Santa Cruz (Fig. 1c),massive ophiolitic rocks that contain basalts ofisland-arc affinity are imbricated with Maastrichtianto Eocene (c. 80–50 Ma) carbonate–arenaceouspre-flysch and flysch (Sepur Formation; Wilson1974; Beccaluva et al. 1995). Aptian and Cenoma-nian (c. 120–95 Ma) sedimentary rocks of thePeten basin of northern Guatemala (Maya Block)record a pelagic facies, reflecting maximum trans-gression and an open marine connection to the Gulfof Mexico to the north, and the Proto-Caribbeanseaway to the south (Fourcade et al. 1994, 1999).During the Late Maastrichtian and Danian (c. 70–60 Ma), the southern Peten basin (Fig. 1b) was adeep siliciclastic sink, the Sepur foredeep basin.The depocentre of this basin shifted from southduring the Late Campanian to north during theLate Maastrichtian and Early Danian; this led tothe demise of the pre-existing carbonate platformover more than 100 km from south to north.

The ophiolites south of the Motagua fault (SouthEl Tambor complex, Fig. 2k) are limited to the northby the Cabanas fault and are overlain by Subinal redbeds and in the south by the Cenozoic–Recentvolcanic arc; they are cut by felsic intrusives.Together with Late Jurassic–Cretaceous ‘melange’,they cover basement of the Las Ovejas complex andthe San Diego Group of the Chortıs Block. Basalts

within these ‘southern’ ophiolites show mid-oceanridge basalt (MORB) affinity (Beccaluva et al.1995). In the north-central Chortıs Block (Fig. 2j),the clastic, partly turbiditic, rift-related Agua FrıaFormation was deformed during Late Jurassic(Viland et al. 1996). It is overlain by the shallow-marine, shelf-type Yojoa Group that contains arcand intra-arc rift deposits (Rogers et al. 2007a).Possibly two molasse sequences are present: themainly red beds of the Albian–Cenomanian Vallede Angeles Group and the ?Eocene–OligoceneSubinal Formation; the latter may have been depos-ited in pull-apart basins along the Motagua faultzone. Thus, molasse-type deposits occur on theChortıs during late Early to Late Cretaceous times(Valle de Angeles Group), whereas marine platformcarbonates developed on the Maya Block (Cobanand Campur Groups). Top-to-north thrust imbrica-tion affected the Jurassic–Cretaceous sequence ofthe Chortıs Block during Late Cretaceous (Donnellyet al. 1990; Colon fold belt of Rogers et al. 2007b).

Basement complexes

Clear definition and subdivision of the rock unitsthat comprise the MSZ (Fig. 1c) remain elusive prin-cipally because Cenozoic faults delineate parts ofthese units (see also Ortega-Gutierrez et al. 2007).This section briefly describes the major units of theMSZ from north (Maya Block) to south (ChortısBlock).

Maya Mountains and Altos Cuchumatanes (MayaBlock). Metamorphic rocks in the Maya Mountainsof Belize (Fig. 1c) comprise low-grade meta-sedimentary rocks with minor rhyolite–dacite inter-calations (Steiner & Walker 1996; Martens et al.2010). Lower Palaeozoic beds are intruded by gran-itoids; Upper Palaeozoic deposits were correlatedwith the Santa Rosa Group based on lithology andCarboniferous–Permian fossils (Bateson 1972).The Altos Cuchumatanes area (Fig. 1c) exposes low-to high-grade basement, intruded by granitoids, andmantled by Santa Rosa Group metasedimentaryrocks (Anderson et al. 1973).

Rabinal complex: Salama schists, San Gabriel unit,and Rabinal granitoids (Maya Block). Van denBoom (1972) called low-grade metasedimentaryrocks south of the Polochic fault zone in centralGuatemala the Salama schists. These rocks weregrouped into an older clastic and volcanic succes-sion, the San Gabriel unit, which is intruded bythe Rabinal granitoids, and a younger clastic andcalcareous succession, probably, part of the SantaRosa Group (Fig. 1c; Donnelly et al. 1990;Ortega-Obregon et al. 2008). The basal conglomer-ate beds of the Santa Rosa Group comprise the

N. AMERICA–CARIBBEAN PLATE BOUNDARY 221

Sacapulas Formation of van den Boom (1972).Ortega-Obregon et al. (2004, 2008) proposed an upto 10 km wide, south-dipping shear zone, the BajaVerapaz shear zone, thrusting (with a minor sinistralstrike–slip component) Chuacus gneiss onto SanGabriel rocks; deformation is most pronouncedwithin 300 m of the Chuacus–San Gabriel bound-ary. The lateral continuation of the shear zone isunknown. The San Gabriel unit has a continentaldepositional environment and is pre-Silurian, givenby the age of the Rabinal granitoids (see below). Itresembles low-grade metasedimentary rocks inwestern Guatemala, which are post-Middle Protero-zoic based on detrital zircon geochronology (Solariet al. 2008). The Santa Rosa Group yielded Early

Carboniferous crinoids (van den Boom 1972) andMississippian (Tournaisian) conodonts (Ortega-Obregon et al. 2008). The Rabinal granitoids arespatially poorly defined, as their sheared marginsand the surrounding sediments have similar appear-ance; they are predominately fault-bounded andwidely tectonized. Here we use the term Rabinalcomplex to describe the in general low-grade rockassemblage that comprises pre-Middle Devonianvolcanoclastic rocks and Ordovician–Early Devo-nian granitoids; it is widely associated with itsclastic–calcareous cover, the Santa Rosa Group. Wenote that there is no obvious difference in Palaeo-zoic lithostratigraphy and magmatism between theMaya Block (Maya mountains, Alto Cuchumatanes)

Fig. 2. (Litho)stratigraphic and pressure–temperature–time–deformation evolution and tectono-stratigraphicinterpretation of the Maya Block, the Motagua suture zone, and the northern Chortıs Block and correlative unitsin southern Mexico and the southern Chortıs Block (see text for discussion of correlations). Ages in italics are detritalU–Pb zircon ages or from inherited zircons in granitoids; other ages are from magmatic rocks. All ages are thosereported in the captions of Fig. 1c and the text. Pressure–time conditions from this paper and literature are reviewedin the text.


and the Rabinal complex, suggesting that thePolochic fault is not a fundamental plate boundaryfault (see Ortega-Gutierrez et al. 2007).

Chuacus complex. In Guatemala, Ortega-Gutierrezet al. (2004) redefined the Chuacus complex asschists and gneisses outcropping between theMotagua fault and the Baja Verapaz shear zone ofthe Salama–Rabinal area (Fig. 1c). Chuacus schistsand gneisses contain quartzþ albiteþwhite micaþepidote/zoisite + garnet + biotite + amphibole +omphacite+ allanite+ chlorite+ calcite+ titanite+rutile (Martens et al. 2005). Banded migmatiteand (garnet) amphibolite (locally relict eclogite),

calcsilicate and marble are uncommon. Local quan-titative petrology suggests upper greenschist- toamphibolite-facies conditions up to 650 8C atc. 1.1 GPa; retrogression to greenschist facies ispoorly documented (van den Boom 1972; Machorro1993). Both foliation concordant and discordantquartz–albite–white mica pegmatites occur. Eclo-gitic relicts contain possible ultra-high pressuremetamorphism, and pyroxene with up to 45 vol%omphacite and garnet together with phengiteþrutileþ zoisiteþ epidote. Thermobarometric ana-lyses yielded c. 740 8C at c. 2.3 GPa with retrogres-sion at c. 590 8C and c. 1.4 GPa. Decompressionseems associated with local partial melting

Fig. 2. (Continued).


(Ortega-Gutierrez et al. 2004). In central-easternGuatemala (San Agustın Acasaguastlan), ortho-gneiss (quartz–monzonite, granodiorite) outcropstogether with high-grade, typical Chuacus para-gneiss; the latter locally shows sillimanite and abun-dant K-feldspar. The volcanosedimentary protolithof the Chuacus complex is probably younger thanc. 440 Ma based on detrital zircon geochronology(Solari et al. 2008).

Existing geochronology: Maya Block, Rabinal andChuacus complexes. U–Pb zircon and monaziteages obtained from two batholiths in the Mayamountains (Steiner & Walker 1996) yielded404–420 Ma (the best constrained samples are418 + 3.6 and 404 + 3.3 Ma); K–Ar biotite agesof these igneous rocks are 230 + 9 Ma (weightedmean of seven dates compiled in Steiner & Walker1996). Martens et al. (2010) discriminated apost-520 Ma (detrital zircon geochronology) volca-nosedimentary sequence (Baldy unit) that is inter-bedded with 404 + 7 Ma rhyolite in its uppersection (Bladen Formation). Granitoids of theAltos Cuchumatanes area yielded 269 + 29 and391.2 + 7.4 Ma (Solari et al. 2008). Weber et al.(2008) reported 482 + 5 Ma U–Pb zircon and473 + 39 Ma Sm–Nd garnet–whole rocks agesfrom S-type granite from the Chiapas Massif ofthe southwestern Maya Block. Three zircon frac-tions of a Rabinal orthogneiss yielded a concordant396 + 10 Ma age (recalculated from the data givenby Gomberg et al. 1968). All zircon fractions of

three Rabinal granites and one pegmatite are discor-dant (Ortega-Obregon et al. 2008); they yieldedlower intercepts of 496 + 26 and 417 + 23 Ma(granite), 462 + 11 Ma (pegmatite), and an upperintercept of 483 + 23 Ma (granite). Muscovitefrom two pegmatites cutting the San Gabriel unityielded K–Ar cooling ages of 453 + 4 and445 + 5 Ma. Ortega-Obregon et al. (2008) sug-gested intrusion of the Rabinal granitoids between462 and 453 Ma. The Matanzas granite of easterncentral Guatemala, possibly intruding the Rabinalcomplex, yielded a two feldspar–whole rock Rb–Sr isochron of c. 227 Ma, and Ar/Ar white micaand biotite ages of c. 212 and 161 Ma, respectively(Donnelly et al. 1990). Four zircon fractions com-bining two Chuacus gneiss samples south ofSalama in central Guatemala gave a lower interceptU–Pb age of 326 + 85 Ma (recalculated fromGomberg et al. 1968). Three single zircon and twosmall populations from refolded leucosome inter-bedded with relict eclogite (El Chol area) yieldedintercept U–Pb ages of 302 + 6 and 1048 +10 Ma (Ortega-Gutierrez et al. 2004). K–Ar andAr/Ar ages from Chuacus-complex rocks aremostly poorly located and documented. Hornblende,white mica and biotite gave Ar/Ar ages of 78–66,76–67 and 67–64 Ma, respectively (Sutter 1979).Amphibole from garnet amphibolite and whitemica from retrograde phengite-bearing gneissyielded K–Ar ages of c. 73 and c. 70 Ma, respect-ively. Coarse- and finer-grained white mica fromfoliated metapegmatite gave c. 73 and c. 62 Ma,



respectively (weighted mean ages from several frac-tions given by Ortega-Gutierrez et al. 2004). Whitemica from Chuacus-complex paragneiss and peg-matite �4 km south of the main trace of the BajaVerapaz shear zone gave ages between 74.1 + 1.3and 65.3 + 1 Ma; they were suggested to dateshearing along the Baja Verapaz zone (Ortega-Obregon et al. 2008). The oldest Ar/Ar age fromthe Chuacus complex is c. 238 Ma (amphibole;Donnelly et al. 1990).

El Tambor complexes: metamorphic–metasomaticrocks of the Motagua suture zone. The El Tamborcomplexes (Fig. 1c) comprise metasedimentaryrocks of a subduction–accretionary complex, sheetsof serpentinized peridotite, serpentinite melangeand exotic blocks, containing eclogite, jadeitite,mylonitized gabbro, (garnet) amphibolite, pillowlavas and radiolarian cherts; as far as is known,the basal contacts are tectonic (e.g. Bertrand &Vuagnat 1975; Harlow 1994; Beccaluva et al.1995; Tsujimori et al. 2005; Chiari et al. 2006).Regional prehnite–pumpellyite-facies metamorph-ism affected the volcanic rocks, forming albiteþchlorite, predominantly pumpellyite, and locallyprehnite. Centimetre- to metre-scale interlayeringof mafic rocks, siliceous marble and quartzite (prob-ably metamorphosed chert) are reported in the LaPita complex of the eastern MSZ (Martens et al.2007); assemblages with barroisiteþ garnetþepidoteþ albiteþwhite mica are characteristic.Jadeitite formed sporadically within the serpentinitematrix of the El Tambor melange; it consists ofjadeitic pyroxene with minor paragonite, phengite,phlogopite, albite, titanite, zircon, apatite andgarnet and formed at 100–400 8C and 0.5–1.1 GPa (Harlow 1994). Two contrasting high-Pbelts contain lawsonite eclogite and zoisite eclogitesouth and north of the Motagua valley, respectively.The southern lawsonite eclogites record unusuallylow-T–high-P with prograde eclogite facies atc. 450 8C and c. 1.8–2.4 GPa, followed by retro-gression with infiltration of fluids at ,300 8C andc. 0.7 GPa (Tsujimori et al. 2004, 2005, 2006).Amphibolitized, paragonite-bearing zoisite eclo-gites in antigorite schist north of the Cabanas faultformed at significantly higher T, 600–650 8C at2.0–2.3 GPa (Tsujimori et al. 2004).

Existing geochronology: El Tambor complexes.Two Sm–Nd mineral isochrons from eclogitessouth of the Motagua fault zone gave ages of131.7 + 1.7 and 132.0 + 4.6 Ma, which are identi-cal to the 130.7 + 6.3 Ma age from eclogitesampled north of the fault (Brueckner et al. 2005).The southern eclogites show MORB major andtrace element patterns and are isotopically depleted(87Sr/86Sr ratios of 0.70374 and 0.70489, 1Nd of

þ8.6 and þ9.2). One northern eclogite has a signi-ficantly more enriched signature (87Sr/86Sr ¼0.70536; 1Nd ¼ 22.1). Harlow et al. (2004)obtained Ar/Ar white mica ages from jadeitite, albi-tite, mica-bearing metamorphic rocks and alteredeclogite within the MSZ, ranging between77–65 Ma for the northern high-P belt and125–113 Ma for the southern belt. In the northernbelt, three phengite-only samples are indistinguish-able within error at 76 Ma, whereas two phengite–phlogopite–paragonite mixtures are younger at 71and 65 Ma. The 600–650 8C peak T of the northernhigh-P rocks suggests that the phengite (closing Tc. 400 8C, von Blanckenburg et al. 1989) dates retro-gression under blueschist-facies conditions whenNa–Al–Si-rich fluids produced jadeitite and albititein serpentinite (Harlow 1994). All dated mineralsfrom the southern belt are phengite; four ages arewithin error at 120 Ma, one is older, two younger.These ages probably date fluid infiltration duringblueschist-facies retrograde metamorphism. Ber-trand et al. (1978) reported K–Ar mineral andwhole-rock ages from poorly located samplesnorth of the Motagua fault zone. We divided theirsamples regionally and recalculated the ages: ametadolerite from the Baja Verapaz ophiolitesheet (Fig. 1c) north of Salama yielded c. 76 Ma,three metadiabase samples from pillow basalt over-lying the Chuacus complex gave a weighted meanwhole-rock age of c. 74 Ma (the 40Ar/36Ar v.40K/36Ar isochron gives atmospheric 40Ar/36Ar),and six amphibole separates from garnet amphibo-lites of the same unit yielded a weighted mean ageof 57.2 + 5.4 Ma (again, atmospheric 40Ar/36Arwith an ‘isochron age’ of c. 56 Ma). All ages prob-ably date metamorphism.

The available petrology (lawsonite and zoisiteeclogites south and north of the Motagua faultzone, respectively), geochronology (c. 120 and76 Ma for blueschist-facies retrogression southand north of the Motagua fault zone, respectively)and isotope geochemistry (isotopically depletedv. enriched south and north of the Motaguafault zone, respectively) demonstrate the existenceof two distinct subduction–accretion belts alongthe MSZ, even though ages of crust formationand eclogite–facies metamorphism may appearsimilar. Herein, we use the terms South and NorthEl Tambor complexes for these two distinctsubduction–accretion belts and their ophioliticsheets, with the South El Tambor complex beinglocated south of the Cabanas strand of theMotagua fault zone, and the North El Tamborcomplex located to its north.

Las Ovejas complex, San Diego phyllite, and Caca-guapa schist (Chortıs Block). The Las Ovejascomplex (Fig. 1c) comprises biotite schist, gneiss,


migmatite, amphibolite, marble and deformedgranitoids. Specifically, amphibolite (�15% of theLas Ovejas complex in Guatemala), garnetiferousgneiss, staurolite schist, garnet–two-mica schistand marble with calcsilicate layers are exposed toge-ther with metaigneous rocks, diorite, tonalite, gran-odiorite and minor gabbro. Mineral assemblagesdefine amphibolite-facies conditions; metamorphicgrade of rocks containing staurolite increasestoward the NW where a sillimanite zone was ident-ified (IGN 1978). Both basement and cover of thenorthern Chortıs Block continue as the BonaccaRidge, forming a series of islands south of theSwan Island Fault (Fig. 1c), the southern boundaryfault to the Cayman Trough. Roatan Island, whichis the best studied (Ave Lallemant & Gordon1999), exposes amphibolite-facies basement with(augen)gneiss, schist, amphibolite (partly metagab-bro), marble and rare pegmatite.

The San Diego metamorphic rocks consist ofphyllite and minor thin quartzite and slate layers(Schwartz 1976). Their stratigraphic age is post-Early Cambrian (youngest detrital zircons arec. 520 Ma; Solari et al. 2008). The Chiquimula bath-olith thermal overprint produced cordierite andandalusite (Clemons 1966). Greenschist-faciesand locally higher-grade metamorphic rocks inHonduras include calcareous phyllite, white mica–chlorite and graphitic schists and slate (‘youngersequence’ of Horne et al. 1976a). The San Diegophyllite of southern Guatemala and northernHonduras has been correlated with the Cacaguapaschists south and east of the Jocotan–Chameleconfault system (Fig. 1c; Burkart 1994). Thegreenschist-facies lithology at Roatan Island com-prises (quartz) phyllite, chlorite schist, serpentinite,chert and marble to limestone; locally granite por-phyry is present (Ave Lallemant & Gordon 1999).

Sanarate complex. We introduce the Sanaratecomplex as a distinct unit separate of the South ElTambor and the Las Ovejas complexes due to its dis-tinct lithology and P–T–d– t parameters; however,its boundaries and lithostratigraphy are mostlyunknown and we may have exaggerated its extendin the geological maps (e.g. Fig. 1c). The complexconstitutes the westernmost basement outcropsouth of the MSZ and comprises partly retrogradeamphibolite, garnet- and quartz-bearing amphi-bolite, retrograde leucogabbro, chlorite–epidote–hornblende schist, micaschist, phyllite and chloritephyllite. Massive amphibolite is characteristic andmost widespread. These lithologies probably rep-resent an oceanic environment. The complex is infault contact with Cenozoic andesite and red bedsand is cut by andesite dykes. Along its easternmargin, the Sanarate-complex rocks shows west-dipping foliation and apparently thrusted top-to-east

over South El Tambor complex metaclastic rocksthat show distinctly lower metamorphic grade.

Existing geochronology: Las Ovejas complex, SanDiego phyllite and Cacaguapa schist, igneousrocks. Most of the available geochronology is, in amodern sense, poorly documented and samples areonly approximately located (Fig. 1c). U–Pb zirconages on igneous rocks are: c. 93 Ma (Rio Juliangranite), c. 80 Ma (sheared Rio Cangrejal orthog-neiss), 81.0 + 0.1 (El Carbon granodiorite),c. 29 Ma (Banderos granodiorite; all Manton &Manton 1984), 124 + 2 Ma (dacite; Drobe &Oliver 1998), and 1017 Ma (orthogneiss; Manton1996). We recalculated the zircon data of the ElCarbon granite of Manton & Manton (1999) andobtained lower and upper intercepts of 30 + 19and 126 + 42 Ma, respectively. Deformed graniticplutons gave a 300 + 6 Ma Rb–Sr whole-rockisochron for the Quebrada Seca complex (Horneet al. 1976a). Granitoids gave the followingRb–Sr ages: 150 + 13 Ma (San Marcos granodior-ite, whole rock; Horne et al. 1976a), c. 69 Ma(K-feldspar and plagioclase, cooling age of shearedEl Carbon granodiorite; Manton & Manton 1984),c. 80 Ma (whole-rock isochron, Trujillo grano-diorite; Manton & Manton 1984), 140 + 15 Ma(Dipilto granite, whole rock; Donnelly et al. 1990),118 + 2 Ma (Carrizal diorite, K-feldspar and plagi-oclase; Manton & Manton 1984), c. 54 Ma (RıoCangrejal granodiorite), 39 Ma (Sula valley (Sulagraben) granodiorite), and 35 Ma (Confadia granite;all biotite cooling ages; Manton & Manton 1984).K–Ar cooling ages of igneous rocks comprise93.3 + 1.9, 80.5 + 1.5 and 73.9 + 1.5 Ma (Telaaugite-tonalite, one amphibole and two biotiteages from different samples; Horne et al. 1974),72.2 + 1.5 and 56.8 + 1.1 Ma (Piedras Negrastonalite, amphibole and biotite; Horne et al. 1974),35.9 + 0.7 Ma (San Pedro Sula granodiorite,biotite; Horne et al. 1974), 122.7 + 2.5 and117.0 + 1.9 Ma (San Ignacio granodiorite, hornble-nde and biotite; Emmet 1983), 114.4 + 1.7 Ma (SanIgnacio adamellite, biotite; Horne et al. 1974),79.6 + 1.6 Ma (gabbro dike, amphibole; Emmet1983), 62.2 + 2.6 Ma (diorite, whole rock; ItalianHydrothermal 1987, reported in Rogers 2003),60.6 + 1.3 and 59.3 + 0.9 Ma (Minas de Oro gran-odiorite, biotite; Horne et al. 1976b; Emmet 1983),58.6 + 0.7 Ma (San Francisco dacite pluton, biotite;Horne et al. 1976b). Ar/Ar ages are from ande-site and diorite of the eastern Chortıs Block(75.6 + 1.3 and 59.9 + 0.5 Ma, amphibole andbiotite; Venable 1994). A K–Ar biotite age of222 + 8 Ma from micaschist probably reflects aTriassic thermal event (MMAJ 1980). Hornblendefrom the amphibolite-facies basement of RoatanIslandyielded36.0 + 1.2 Ma,andagraniteporphyry


from Barbareta Island, east of Roatan Island, yieldeda 39.4 + 2.8 Ma zircon fission-track age (closing Tc. 250 8C; Ave Lallemant & Gordon 1999).

New geochronology

ID-TIMS and ion-probe U–Pb geochronology canbe found in the Supplementary materials, andFigs 3 & 4 show the U–Pb data in Concordia dia-grams and present representative zircon CL pictures,respectively. Our new Ar/Ar and Rb–Sr mineralages are summarized in Table 1 and the Supple-mentary material, respectively, and are plotted inFig. 5. In this paper, we discuss Phanerozoic U–Pb ages; late Middle Proterozoic ages, which wefind in both the Maya and Chortıs Blocks, will bereported elsewhere. Our apatite fission-track(AFT) and titanite fission-track (TFT) data arelisted in the Supplementary material; confinedtrack-length data appear in the Supplementarymaterial. Figure 6 depicts AFT age v. elevation dia-grams for the Rabinal–Chuacus and Las Ovejascomplexes of Guatemala/Honduras and the ChiapasMassif of southeastern Mexico (Fig. 1c), and mod-elled T– t paths for selected samples (Chuacus andLas Ovejas complexes, granitoids). We excludedthe samples from western Guatemala from the agev. elevation data (too few readings and widelyspaced) and also emphasize that the plotted datado not strictly fulfill the criteria for exhumationestimates from age v. elevation data; for example,the samples cover considerable horizontal distances.We assigned errors to the elevation data, as severalsamples were collected in the early-1990s whenhand-held GPS-receiver precision was still inferiorand large-scale topographic maps were locallyunavailable. We emphasize that the regressiondata that we present are first-order estimates.Figure 7 shows the regional distribution of ages,highlights age groups, and depicts T– t fields forselected regions. We employed standard closingtemperatures for the T– t diagrams: Ar hornblende,high-T steps, 525 + 25 8C, low-T steps, 400 and300 + 100 8C, depending on grain size; Ar whitemica, 350 + 25 8C; Ar biotite, 300 + 25 8C; Ar K-feldspar, 400 + 100 8C for high-T steps and200 + 25 8C for low-T steps, depending on thelikely presence of single or multiple domains; Rbwhite mica, 500 + 25 8C, except for centimetre-sized, pegmatitic white mica, 600 + 100 8C; Rbbiotite, 275 + 25 8C; TFT, 275 + 50 8C; AFT,100 + 25 8C; exceptionally, errors were sethigher, imposed by variable grain size and strongdeformation. Zircon crystallization was taken as800 + 50 8C, except where zircon growth occurredduring peak-T metamorphism and related migmati-zation, which is constrained by our petrology (seebelow). We excluded a few samples from the T– t

diagrams: ages from sample 5C-36 of the Chuacuscomplex immediately below North El Tamborcomplex serpentinite, as these rocks show extreme,late-stage metasomatism; ages from very coarse-grained, pegmatitic white mica. In the followingsection, we present the data; we relate the ages tothe main temperature–deformation (T–d) eventsin the petrology and structure sections.

U–Pb zircon geochronology

Rabinal complex. Sample G18s (stop G1; Figs 3a &7a) is a migmatitic gneiss from western Guatemalaoverprinted by lower greenschist-facies defor-mation. Four fractions with identical 75–125 mmgrains but different magnetic properties were uti-lized for ID-TIMS analysis. A discordia fit to threefractions yields intercepts of 220.0 + 3.5 and1292 + 39 Ma (MSWD ¼ 0.78); one of these frac-tions is concordant at 215.92 + 0.18 Ma. The fourthfraction is concordant at 268.04 + 0.55 Ma. CLimages (Fig. 4) depict well-rounded cores withvariable but mostly high U-content, and low-U (onaverage 63 ppm), oscillatory zoned, idiomorphicrims with Th–U ratios (on average 0.7) that suggestmagmatic origin. The weighted average of sixgrains analysed with the SHRIMP gives a207-corrected 206Pb–238U age of 217.7 + 7.3 Ma,within error of the concordant TIMS fraction. Weinterpret this age as the age of leucosome crystalli-zation within the western Rabinal complex; thewell-rounded cores suggest a paragneiss protolithand the concordant fraction at c. 268 Ma may indi-cate an earlier phase of crystallization. SampleG27s (stop G57, western Guatemala, Figs 3a &7a) is orthogneiss that shows low-T mylonitizationalong discrete sinistral shear zones. Five fractionsvarying in grain size or magnetic properties are alldiscordant but define a chord with lower and upperintercept ages at 284 + 58 and 958.7 + 6.8 Ma,respectively (MSWD ¼ 0.63); this rock is probablya Permian intrusion. Sample GM8s (stop G8;Figs 3a & 7a) is granite overprinted by greenschist-facies mylonitization from western Guatemala. Fourzircon fractions with variable grain size and mag-netic properties define lower and upper interceptsat 165.2 + 5.4 and c. 990 Ma, respectively (ID-TIMS); one fraction is weakly reverse discordantand two fractions define a concordant age of163.80 + 0.38 Ma (Fig. 3a). CL images show oscil-latory zoned rims around Proterozoic cores (Fig. 4)and oscillatory zoned coreless grains with markedchanges in luminescence from centre to rims; theages of these grains are identical within error. Con-sequently, the Th–U ratios obtained from SHRIMPspots are variable and are with one exception .0.2,implying magmatic origin. The 207-corrected agesfor a group of six rims and coreless grains yield


Fig. 3. U–Pb zircon data and age interpretation. See the Supplementary material for sample location and data listings. ID-TIMS results are shown with 2s error ellipses anddiscordia are given as lines. SHRIMP results show 2s error ellipses and are corrected for common Pb. Selected age groups are displayed as weighted average and relative probabilitydiagrams (206Pb–238U ages are 207-corrected). Data evaluation and plotting was supported by the program package of Ludwig (2003). MSWD, mean square weighted deviation. Theupper intercept ages and ages older than Early Palaeozoic will be discussed elsewhere. (a) Data from the Rabinal complex of the Maya Block.

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Fig. 3. (Continued) (b) Data from the Chuacus complex.

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Fig. 3. (Continued) (c) Data from the Chortıs Block; all SHRIMP ages are from the northeastern edge of the Las Ovejas complex and granitoids intruding it. Th–U diagrams aiddiscrimination of metamorphic and magmatic zircons (or rims).

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Fig. 3. (Continued) (d) Data from the Chortis Block, TIMS U–Pb zircon ages.

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a weighted average of 171.0 + 2.1 Ma. Graniteemplacement is either given by this date or isslightly younger (concordant TIMS fractions).

Sample 5CO-1 is two-mica, K-feldspar–porphyroblastic orthogneiss, mylonitized during

low-T overprint; it defines the ‘Rabinal’ granitoidsin the area of San Miguel Chicaj, central Guatemala(Fig. 7a). Rounded, often mottled, Proterozoiccores are surrounded by in general thin, structure-less, often high-U rims that have variable but

Fig. 4. Cathodoluminescence images of representative zircons. Numbers give sample, spot and spot age (in Ma, belowor above the sample number).


Table 1. 40Ar/39Ar data

Sample Stop Mineral J Weight(mg)

Grain size(mm)

Total fusionage (Ma)

Weightedmean age

(Ma)

Isochron age(Ma)

MSWD 40Ar/36Ar %39Arused

Steps

Rabinal complexG19s G1 Phe 0.003192 10.1 10–20 66.66 + 0.33 66.53 + 0.33 66.50 + 0.37 4.1/2.0 296 + 4 61 6–16/21

Phe 0.003185 10.1 20–40 67.92 + 0.34 67.38 + 0.33 67.00 + 0.37 4.1/1.9 307 + 5 79 7–19/23G27s G57 Bt 0.004555 0.3 250 37.38 + 1.37 35.06 + 1.29 37.21 + 10.06 0.6/2.4 350 + 74 48 5–10/15

Chuacus complexG10s G29 WM 0.004564 1.2 100–200 66.04 + 0.65 66.00 + 0.65 65.96 + 0.66 14.6/1.7 299 + 9 93 3–21/26G12s G31 WM 0.0007854 n.d. n.d. 71.80 + 0.89 72.59 + 0.65 67.8 + 5.0 31 297 + 64 100 1–19/195CO-4b 5CO-4 Bt 0.004480 4.3 250 64.31 + 0.63 64.34 + 0.63 64.80 + 0.68 26.6/1.8 264 + 11 98 2–17/19

WM 0.004590 2.9 250 67.77 + 0.67 67.51 + 0.66 67.56 + 0.73 174/1.7 286 + 17 100 1–17/175CO-5a 5CO-5 WM 0.004586 2.1 250 234.4 + 2.2 225.1 + 2.1 225.2 + 2.2 0.9/2.6 294 + 3 5 1–5/23

Hbl 0.004581 8.9 250 670.61 + 5.60 — — — — 100 1–20/205PB-9b 5PB-9 WM 0.004570 5.1 250 60.27 + 0.59 — — — — 100 1–27/275PB-13b 5PB-13 Hbl 0.0007852 n.d. n.d. 94 + 18 69.3 + 1.4 64.5 + 5.4 0.31 1283 + 270 35 5 of 125PB-14 5PB-14 WM 0.004568 2.9 250 64.67 + 0.63 65.01 + 0.64 64.93 + 0.64 1.2/2.3 297 + 1 68 14–20/20854F 854 Bt 0.014640 28 250 125.3 + 3.5 — — — — 100 10/10880C 880 WM 0.014640 16.8 250 102.4 + 2.7 101.0 + 2.5 100.4 + 1.5 0.9/2.6 335 + 32 88 5–9/10883B 883 WM 0.014640 17.1 250 75.2 + 1.8 74.7 + 1.5 74.2 + 0.8 0.6/2.4 338 + 21 93 5–10/10883B 883 Bt 0.014640 14.9 250 68.2 + 2.0 68.6 + 2.0 71.3 + 1.8 9.0/3.0 146 + 90 100 1–8/8887C 887 WM 0.014376 5.4 250 70.5 + 1.2 69.8 + 1.2 70.9 + 3.8 13 321 + 120 78 1–7/10

Metasomatized southern Chuacus complexG6s G26 WM 0.004565 0.6 c. 63 71.40+ 0.80 62.01 + 0.69 59.55 + 2.37 2.2/2.3 497 + 367 48 6–12/16G7s G27 WR rhyolite 0.004446 20.0 n.a. 68.37 + 0.94 69.91 + 1.06 63.66 + 9.19 9.5/3.0 305 + 15 59 2–5/75C-36a 5C-36 WM 0.004579 1.9 250 48.04 + 0.47 48.32 + 0.48 48.20 + 0.50 16.7/1.8 298 + 3 86 9–22/22

Sanarate complexG3s G23 WM 0.004534 2.2 c. 63 168.05 + 1.61 170.56 + 1.63 151.2 + 13.22 80/2.4 1339 + 918 69 3–8/17G5s G24 Hbl 0.004519 9.5 250 128.08 + 1.29 155.85 + 1.50 155.92 + 1.91 75/2.7 291 + 8 36 13–17/17

(Continued)

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Table 1. Continued

Sample Stop Mineral J Weight(mg)

Grain size(mm)

Total fusionage (Ma)

Weightedmean age

(Ma)

Isochron age(Ma)

MSWD 40Ar/36Ar %39Arused

Steps

Western Las Ovejas complex, Guatemala5C-2c 5C-2 Hbl 0.000786 n.d. n.d. 22.1 + 6.6 19.2 + 1.5 20.1 + 6.4 0.32 280 + 36 73 1–4/65C-5 5C-5 Bt 0.004574 0.5 c. 63 29.21 + 0.29 29.18 + 0.29 29.90 + 0.33 3.4/2.6 255 + 6 88 2–6/75C-23d 5C-23 Hbl 0.004609 23.4 250 35.99 + 0.37 35.43 + 0.36 35.81 + 0.40 1.6/2.0 276 + 9 64 6–15/175C-23c 5C-23 Bt 0.004612 0.8 250 25.57 + 0.26 25.24 + 0.25 25.34 + 0.26 0.3/3.0 291 + 2 60 3–6/95C-26c 5C-26 Hbl low-T 0.004607 18.9 250 27.84 + 0.28 17.76 + 0.18 18.21 + 0.45 48.3/2.6 272 + 18 45 1–5/19

medium-T 21.73 + 0.24 21.72 + 0.35 2.4/2.4 384 + 3 13 10–15/19high-T 38.01 + 0.38 35.78 + 2.28 43.8/3.8 314 + 18 17 15–17/19Bt 0.004605 2.9 250 13.84 + 0.62 20.49 + 0.24 21.08 + 0.84 18.1/2.0 294 + 2 76 4–9/11

5C-26d 5C-26 Kfs 0.004600 12.7 400 30.25 + 0.59 30.42 + 0.32 29.94 + 0.54 6.4/2.3 297 + 1 89 10–16/175C-33 5C-33 Bt 0.004597 2.5 250 22.42 + 0.24 23.04 + 0.23 22.80 + 0.26 2.2/2.1 297 + 1 58 6–13/17

Kfs low-T 0.004524 26.8 400 129.9 + 2.0 7.4 + 1.1 6.7 + 3.2 1.7/2.6 298 + 25 23 6–10/21medium-T 31.38 + 2.78 32.98 + 10.47 0.1/3.8 292 + 7 7 11–13/21

5C-37b 5C-37 Hbl high-T 0.004566 8.6 250 41.06 + 0.43 39.11 + 0.55 39.10 + 3.93 1.2/2.1 512 + 117 37 8–11/12low-T 18.23 + 0.53 18.27 + 1.63 1.7/2.3 362 + 20 20 1–3/12

5C-37d 5C-37 Bt 0.004593 2.9 250 17.81 + 0.18 18.15 + 0.19 19.49 + 1.18 60/2.3 289 + 6 74 4–10/125PB-5a 5PB-5 Hbl 0.0007864 n.d. 250 30.4 + 1.0 31.03 + 0.70 29.9 + 1.4 67/– 243 + 77 96 2–13/15

Bt 0.004572 1.6 250 28.22 + 0.28 28.20 + 0.28 28.29 + 0.32 18.2/2.1 289 + 6 100 1–8/8

Eastern Las Ovejas complex, Honduras5H-3a 5H-3 Bt (laser) 0.004420 s.g. 250 25.58 + 0.26 18.98 + 0.22 19.00 + 2.78 17.4/2.1 661 + 412 100 1–10/105H-4b 5H-4 Bt (laser) 0.004424 s.g. 250 25.58 + 0.34 25.70 + 0.34 23.05 + 1.86 1.6/3.0 314 + 13 100 1–4/4

Hbl 0.004532 5.3 250 36.87 + 0.40 38.81 + 0.40 38.86 + 1.78 93.7/2.6 291 + 135 92 5–9/105H-4c 5H-4 Bt 0.004577 2.3 250 24.77 + 0.25 25.75 + 0.26 24.67 + 0.44 0.6/3.0 311 + 4 54 5–8/9

Hbl 0.004527 7.5 250 32.57 + 0.40 31.53 + 0.32 31.44 + 0.65 52.3/2.1 297 + 6 98 4–11/115H-4d 5H-4 Hbl 0.0007881 n.d. n.d. 28.6 + 2.0 30.0 + 1.5 26.1 + 3.6 6.1/– 334 + 16 100 1–17/17

Granitoids of the arc complex5C-11 5C-11 Btþ Chl 0.004422 s.g. n.a. 84.51 + 0.89 89.98 + 0.89 — — — 100 1–10/105C-21 5C-21 Kfs (laser) 0.004416 s.g. 400 20.03 + 0.23 19.09 + 0.25 18.80 + 0.45 0.3/3.0 297 + 1 58 5–9/95H-5 5H-5 Kfs low-T 0.004529 12.3 400 32.98 + 0.33 28.29 + 0.28 28.31 + 0.36 2.4/2.6 291 + 59 27 10–14/335H-13 5H-13 Bt (laser) 0.004425 3.0 250 57.71 + 0.70 60.89 + 0.61 61.95 + 0.86 9.7/2.4 265 + 14 100 1–6/6

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Southern MexicoML-18 ML-18 Bt 0.004540 1.1 250 148.82 + 1.43 162.2 + 0.8 — — — 10 19–21/21ML-37 ML-37 WM high-T 0.004542 1.7 250 302.76 + 2.76 315.09 + 2.87 316.3 + 0.4 5/3 289 + 2 37 16–20/20ML-37 ML-37 WM low-T 0.004542 1.0 250 302.76 + 2.76 246.70 + 2.76 255.39 + 22.70 68/3.8 289 + 14 2 1–3/20MU-12 MU-12 Bt 0.0015255 1.0 250 29.69 + 0.35 28.60 + 0.29 28.62 + 0.40 0.02/3.8 295 + 2 58 13–16/17MU-13 MU-13 Bt 0.0015249 1.2 250 35.06 + 1.24 34.38 + 0.39 35.12 + 0.70 0.5/3.0 275 + 12 56 11–15/16ML-39 ML-39 Bt 0.0015307 1.5 250 15.22 + 0.46 14.57 + 0.59 13.81 + 1.90 1.6/2.6 300 + 10 51 1–5/15PA3-4 PA3-4 Hbl 0.0042502 s.g. — 8.1 + 0.6 8.8 + 1.5 1.9 290 + 27 100 1–5/5PA3-4 PA3-4 Bt 0.0042502 203/s.g. 3 exp. — 8.5 + 0.3 2.1 Forced 100 1–9/9CB35 CB35 Hbl 0.0036469 198/320 2 exp. — 28.4 + 1.3 3.0 293 + 25 — 1–7/9CB35 CB35 Bt 0.0035870 35/203 2 exp. — 8.2 + 0.1 0.57 297 + 3 100 1–12/12CB31 CB31 Hbl 0.0049746 207/305 2 exp. — 15.5 + 1.2 0.89 305 + 19 — 1–6/8CB34 CB34 Hbl 0.0035870 195/306 2 exp. — 11.0 + 1.8 0.33 307 + 6 100 1–9/9PI28-1 PI28-1 Hbl 0.0042502 s.g. — 8.1 + 0.3 — — — 100 1–4/4PI28-1 PI28-1 Bt 0.0042502 191/s.g. 3 exp. — 8.0 + 0.5 1.3 303 + 3 100 1–8/8

J is the irradiation parameter; MSWD is the mean square weighted deviation, which expresses the goodness of fit of the isochron; isochron and weighted mean ages are based on fraction of 39Ar and steps listedin the last two columns. WM(P)A is the weighted mean (plateau) age. Abbreviations: sg, single-grains (laser total-fusion ages); Phe, phengite; WR, whole rock; WM, white mica; Hbl, amphibole; Kfs,potassium feldspar; Bt, biotite; Chl, chlorite; low-T, low-temperature steps; medium-T, intermediate-temperature steps; high-T, high-temperature steps; exp., heating or laser degassing experiments on differentfractions of minerals from the same sample; n.a., not applicable; n.d., not determined. Errors are 2s.Age interpretation – Rabinal complex: G19s fine-grained phengite: unified age for two grain-size fractions: 67 + 1 Ma; G27s biotite: WMA calculated for 40Ar/36Ar ¼ 350; 37 + 5 Ma.Central Chuacus complex: 5CO-4b biotite: 64.5 + 1.0 Ma; 5CO-4b white mica: 67.5 + 2.0 Ma. G10s white mica: 66 + 1 Ma; G12s white mica: 70 + 3 Ma; 5CO-5a white mica: hump-shaped spectrumindicates 40Ar excess, 219–243 Ma; low-T plateau at 225 + 5 Ma; 5CO-5a hornblende: loss profile from c. 720 to c. 153 Ma; 5PB-9b sericitic white mica: 52–62 Ma hump-shaped spectrum indicates40Ar excess; 5PB-13b hornblende: mixing of different age components; minimal age of old component 137 Ma, maximal age of young component 20 Ma; major age component at 65 + 6 Ma; 5PB-14white mica: loss profile from c. 65 to c. 50 Ma; mid and high-T ‘plateau’ at 65.2 + 0.8 Ma; 854F biotite: age range from 148 to 94 Ma; 880C white mica: age range from 127 to 100 Ma; likely Cretaceousmetamorphic overprint on Early Mesozoic or older metamorphic rock; 883B white mica: 75 + 2 Ma; 883B biotite: 68.6 + 2.0 Ma; 887C white mica: 70.0 + 1.5 Ma.Metasomatized southern Chuacus complex: G6s white mica: U-shaped spectrum indicates 40Ar excess; 60 + 5 Ma; G7s rhyolite gash filling: 67 + 5 Ma; 5C-36a white mica: 48 + 1 Ma.Sanarate complex: G3s white mica: loss profile from c. 195 to c. 137 Ma; mid-T ‘plateau’ at 155 + 10 Ma; G5s hornblende: loss profile from c. 157 to c. 45 Ma; high-T ‘plateau’ at 156 + 5 Ma, low-T ‘plateau’at 60 + 20 Ma.Western Las Ovejas complex, Guatemala: 5C-2c hornblende: 20 + 2 Ma; 5C-5 biotite: loss profile from c. 31 to c. 26 Ma; mid-T ‘plateau’ at 29 + 2 Ma; 5C-23c biotite: 25.3 + 0.5 Ma; 5C-23d hornblende:35.5 + 1.0 Ma; 5C-26c hornblende: loss profile from c. 64 to 17 Ma; low-T ‘plateau’ at 18 + 2 Ma, intermediate-T ‘plateau’ at 21 + 2 Ma and high-T ‘plateau’ at 36 + 3 Ma; intermediate and low-T ‘plateau’and chlorite-biotite ages are identical within error at 20 + 2 Ma; 5C-26c chlorite-biotite mixture: loss profile from c. 28 to c. 3 Ma; mid-T ‘plateau’ at 20 + 2 Ma; 5C-26d potassium-feldspar: 30 + 1 Ma;5C-33 biotite: loss profile from c. 26 to c. 15 Ma; mid-T ‘plateau’ at 23 + 2 Ma; 5C-33 potassium-feldspar: loss profile from c. 370 to c. 6 Ma, low-T 40Ar excess, fast cooling at 32+4 and 7 + 2 Ma; 5C-37bhornblende: high-T WMA calculated for 40Ar/36Ar ¼ 553; 39 + 3 Ma; low-T WMA calculated for 40Ar/36Ar ¼ 362; 18 + 3 Ma; 5C-37d biotite: loss profile from c. 26 to c. 9 Ma; mid-T ‘plateau’ at19 + 2 Ma; 5PB-5a hornblende: 30 + 2 Ma; 5PB-5a biotite: 28.2 + 0.4 Ma.Eastern Las Ovejas complex, Honduras: 5H-3a biotite: single-grain laser-fusion age (10 grains); WMA calculated from 40Ar/36Ar ¼ 661; 22 + 4 Ma; 5H-4b biotite: single-grain laser-fusion age (4 grains) at25 + 3 Ma; 5H-4c biotite: loss profile from c. 28 to c. 18 Ma; minor 40Ar excess; mid-T ‘plateau’ at 25 + 2 Ma; Integrated biotite age of stations 5H-3 and 5H-4 at 24 + 2 Ma; 5H-4b hornblende: 38 + 2 Ma;5H-4c hornblende: 31 + 5 Ma; 5H-4d hornblende: 29 + 3 Ma; Integrated hornblende age of stations 5H-4 at 34 + 4 Ma.Granitoids: 5C-11 chloritized biotite: single-grain laser-fusion age (10 grains) at 90 + 10 Ma; 5C-21 potassium-feldspar: single-grain laser-fusion age (8 grains) at 20 + 1 Ma; 5H-5 K-feldspar: loss profilefrom c. 37 to c. 28 Ma, low-T ‘plateau’ indicates fast cooling at 28.3 + 0.5 Ma; 5H-13 biotite: single-grain laser-fusion age (6 grains) at 59 + 3 Ma.Southern Mexico: ML-18 biotite: loss profile from 163 to 88 Ma, high-T ‘plateau’ at 162 + 2 Ma, low-T ‘steps’ at 90 + 2 Ma; ML-37 white mica: loss profile from 316 to 238 Ma, high-T ‘plateau’ at315 + 3 Ma, low-T ‘plateau’ at 247 + 3 Ma; MU-12 biotite: 29 + 1 Ma; MU-13 biotite: 35 + 1 Ma; ML-39 biotite: 15 + 2 Ma; PA3–4 hornblende: 8.5 + 1.5 Ma; PA3–4 biotite: 8.5 + 1.0 Ma;CB35 hornblende: 28.5 + 1.5 Ma; CB35 biotite: 8.2 + 0.5 Ma; CB31 hornblende: 15.5 + 1.5 Ma; CB34 hornblende: 11.2 Ma; PA28-1 hornblende: 8.1 + 0.5 Ma; PI28-1 biotite: 8.0 + 0.5 Ma.

N.

AM

ER

ICA

–C

AR

IBB

EA

NP

LA

TE

BO

UN

DA

RY

235

Fig. 5. New 40Ar/39Ar mineral spectra and Rb–Sr whole rock–feldspar–mica ‘isochrons’. See the Supplementarymaterial for sample locations, age data and interpretations. Weighted mean ages (WMA) were calculated usingblack steps of the given temperature range; TFA, total fusion age; IA, isochron age. Uncertainties are +1s.


Fig. 5. (Continued) ‘Atm.’ in the isochron diagrams is the 40Ar/36Ar ratio of the atmosphere (295.5). MSWD, meansquare weighted deviation. Abbreviations: wr, whole rock; Pl, plagioclase; Kfs, K-feldspar; WM, white mica;Bt, biotite.










magmatic Th–U ratios (mean 0.4) (Fig. 4). The207-corrected spot ages, mostly from rims but alsofrom a few coreless grains with relatively flat CLresponse, range from 410 to 485 Ma (Fig. 3a).There may be two age groups with weightedaverage ages of 464.6 + 5.3 and 412.9 + 3.7 Ma;Th–U ratios are similar. We suggest that theseages indicate prolonged Ordovician–Early Devo-nian magmatism (see below).

Sample GM1 (stop GM15, central Guatemala;Figs 3a & 7a) is quartz phyllite close to Salama(central Guatemala); these rocks were mapped as

Santa Rosa Group. Zircons are rounded but a sig-nificant portion is sub- to euhedral. Besides ages.1 Ga, we obtained four concordant ages at416.5+1.6, 282.2+1.7, 260.4+ 2.9 and 245.1 +2.9 Ma. The Devonian spot has low Th–U (0.012),the other zircons (Th–U on average 0.64) probablycrystallized from a magma (note difference in lumi-nescence of 416 and 257 Ma grains in Fig. 4). Ourages imply post-Early Triassic deposition for thesemetasediments.

Samples GM13s (stop GM42) and GM14s (stopGM43) (Figs 3a & 7a) are from western Guatemala.



GM13s is (?ortho)gneiss and GM14s granite intrud-ing gneisses along the Polochic fault. Seven zirconfractions of GM13s with different magnetic proper-ties and, in part, grain size yielded lower and upperintercepts at 20 + 15 and 941 + 7.9 Ma, respect-ively. We analysed six fractions of GM14s differingin grain size and magnetic properties. All are discor-dant and yield lower and upper intercepts at 21 + 33and c. 894 Ma; five fractions yield a lower interceptat 13 + 29 Ma. Phanerozoic spot ages of zirconsfrom GM14s gave one concordia age at 425.5 +6.0 Ma and 14 spots with a 14.65 + 0.42 Maweighted average age. These ages demonstrateMiddle Miocene Central American arc magmatism.

Chuacus complex. Sample 5CO-4b, high-P meta-morphic (see below) garnet orthogneiss from thenorthern Chuacus complex, central Guatemala(Fig. 7a), comprises complex, mostly Proterozoiczircons. Zircon grain shapes include both roundedand prismatic–idiomorphic. In general, low-Ucores are surrounded by variable, higher-U rimsthat both have Proterozoic ages (1180–1017 Maconcordant spots; Fig. 3a). Spots that include theformer grain portions and tiny, very luminescentlow-U outer rims, irresolvable with the spot sizewe employed, yielded the youngest ages. Thelower intercepts of discordia fitted to two groupsof grains indicate crystallization at 638–477 Ma(numbers include errors), predating the high-Pmetamorphic event. Sample 5CO-6b (south-centralChuacus complex; Fig. 7a), a two-mica orthogneiss,is intruded by very coarse-grained pegmatite andinterlayered with paragneiss and garnet amphibo-lite; the latter contains relict eclogite (see petrologyfor analysis of equivalent eclogite sample 5CO-5a).All rock types are ductilely deformed. The amphi-bolite forms boudins in the orthogneiss and ismore strongly deformed. We interpret the massiveorthogneiss as a sill and the eclogite-facies meta-morphism pre-dating its emplacement. The zirconsof the orthogneiss are complex (Figs 3a & 4): wedetected a few well-rounded Proterozoic cores,one core that has flat CL response, low, likelymetamorphic Th–U (0.020) and a 207-correctedage of 402.9 + 4.0 Ma, and many oscillatory zonedgrains with magmatic Th–U (on average 0.82);three of the latter spots give a concordia age of238.0 + 2.5 Ma, and eight spots yield a well-defined group with a weighted average of 218.2 +2.3 Ma. Almost all grains have mostly thin(5–40 mm) overgrowths with flat CL response andlow Th–U (on average 0.01) that span 89 to 69 Ma.The oldest grain has the highest Th–U and probablysamples different age domains (core and rim).Selected omission of grains from the remaininggroup yields one statistically significant group ofsix grains with a weighted average age of

74.24 + 0.52 Ma and five younger spots that mayconstitute a cluster at 70.3 + 1.5 Ma; the lattergroup may be geologically meaningless, if it reflectsprogressive Pb loss of grains from the former group.Sample 5CO-6b is most straightforwardly inter-preted as a Triassic magmatite that inherited Proter-ozoic and Devonian zircons and experienced high-Tmetamorphism during Late Cretaceous, probablycoeval with pegmatite injection; similar verycoarse-grained pegmatites also yielded Late Cretac-eous K–Ar (Ortega-Gutierrez et al. 2004) andRb–Sr white mica cooling ages (see below). Theregional significance of magmatism is supportedby Triassic hornblende and white-mica coolingages from this region (238–225 Ma, Donnellyet al. 1990; see below). As we interpret the ortho-gneiss as a sill, the high-P event is probably pre-Middle Triassic.

Chortıs Block. Sample 5H-3 is syntectonic, syn-migmatitic, plagioclase-rich granite gneiss associ-ated with migmatite, leucogranite and gabbro–diorite; it is typical of the Las Ovejas complex andcrops out at the western flank of the Sula graben innorthwestern Honduras (Fig. 7a). The 207-correctedspot ages on typically magmatic zircons (Figs 3c &4) range from 40.0 + 0.4 to 36.5 + 0.9 Ma withthree grains defining a concordia age at37.0 + 0.52 Ma. We interpret this age as the timeof magmatism and deformation. Sample 5H-9acomprises orthogneiss and migmatite south of theJocotan fault zone (Fig. 7a) that are distinctlydifferent in grain-size (coarse) and deformation(weak) from the gneisses and migmatites of theLas Ovejas complex along the MSZ. Besides Proter-ozoic ages, there are concordant spot ages at400.3 + 3.5 and 328.9 + 4.0 Ma. There are twomain age groups: 272.8 + 2.8 Ma (eight spots)and 244.8 + 2.3 Ma (four spots). These groupshave distinctly different Th–U ratios andCL-response (Figs 3c & 4). The younger group,probably metamorphic, has mean Th–U at 0.04and appears dark and unzoned in CL. The oldergroup has mean Th–U of 0.7 and clear oscillatoryzoning. We suggest Early Permian protolith crystal-lization and Early Triassic metamorphism. Sample5H-11a is low-T, mylonitic orthogneiss (phyllonite),probably metarhyolite intercalated with paragneissand marble; the sample is also from the Jocotanfault zone (Fig. 7a). Six out of seven 207-correctedspot ages yield a 130.5 + 2.7 Ma concordia age(Fig. 3c); Th–U ratios are on average 0.41. Weinterpret this date as reflecting Early Cretaceousmagmatism. Sample 5PB-5c is a small orthogneisspod in massive migmatite that is enclosed in largegranite in easternmost Guatemala (Fig. 7a); it isassociated with garnet amphibolite (see petrology).Three spots define a 36.2 + 2.1 Ma concordia age


Fig. 6. Low-temperature thermochronology data. (a) Apatite fission-track (AFT) age v. elevation plots for the centralChuacus complex (top), the Las Ovejas complex, and undeformed arc granitoids. Locations and fission-track parametersare given in the Supplementary material. (b) AFT temperature–time (T– t) paths for rocks of the Rabinal (westernGuatemala), Chuacus (central Guatemala) and Las Ovejas (western area) complexes, and three arc granitoids; the latterare: 5C-6, Chiquimula batholith in the Las Ovejas complex; 5C-32, small granite east of the Chiquimula batholith;GM15s, augengneiss of the westernmost Rabinal complex overprinted by c. 15 Ma granite. Good-fit solutions


(Fig. 3c), probably the age of migmatization. One ofthe spots is from a relatively high-U, oscillatoryzoned central grain portion that is surrounded by arim poorer in U (Fig. 4); this rim is c. 5 Ma(31.2 + 1.2 Ma) younger. The major group ofPhanerozoic ages spans 140–191 Ma with two,probably significant, age groups: 158.8 + 0.52 Ma(n ¼ 4) and 165.79 + 0.87 Ma (n ¼ 13); the con-sistently lower Th–U ratios (on average 0.12) ofthe former (the latter has on average 0.36) supportthe grouping. There are other Phanerozoic con-cordant spots at 139.8 + 1.0 Ma (n ¼ 2),190.6 + 1.4 Ma (n ¼ 1) and 240.6 + 3.5 Ma(n ¼ 1). We suggest a Middle Jurassic granite proto-lith that inherited Proterozoic to Early Jurassiczircons. Sample 92RN412 is granitic gneiss sur-rounded by amphibolite-facies rocks from the LasOvejas complex of northwestern Honduras(Fig. 7a); it is grossly the same rock as 5H-3. Rb–Sr and K–Ar biotite cooling ages are c. 39 andc. 36 Ma, respectively (Horne et al. 1974, 1976a).Three of four zircon fractions with different grainsize and magnetic properties plot close to concordia;one fraction is concordant at 38.3 + 0.53 Ma. Allfractions define intercepts at 36.6 + 3.1 and c.1014 Ma (Fig. 3c); three fractions are at37.49 + 0.53 and c. 1136 Ma. The c. 38 Ma age iswithin error identical to sample 5H-3. The age con-strains deformation within the eastern Las Ovejascomplex to be younger than 39 Ma (see below).Sample 94SA602 is amphibolite within the meta-morphic sequence that contains the granite gneissof sample 92RN412; the protolith was probablygabbro. We obtained three zircon fractions thatplot close to and on concordia, defining a lowerintercept at 88 + 51 Ma (Fig. 3c); the concordantfraction is at 88.27 + 0.56 Ma, probably datinggabbro crystallization. Sample 93LP522 is granitedirectly south of the Jocotan–Chamelecon faultzone (Fig. 7a). Four zircon fractions define a lowerintercept at 167.6 + 2.6 Ma (MSWD ¼ 10.9), onconcordant fraction is at 167.99 + 0.46 Ma(Fig. 3c); the c. 168 Ma age is within error identicalto sample 5PB-5c and demonstrates Middle Jurassicmagmatism in the Chortıs Block. Samples SF1D,92QO404 and 86TA100 typify orthogneissesembedded in the Cacaguapa schists of the central

Chortıs Block in central Honduras (Fig. 1c). SF1Dis weakly deformed granite gneiss surrounded byJurassic metasedimentary rocks of the Agua FrıaFormation; the contact relationships are unclear.All five fractions are discordant and lower interceptages between 275 and 234 Ma can be defined(Fig. 3c). 92QO404 is mega-crystic, K-feldsparmetagranite surrounded by phyllite. Five out of sixfractions define intercepts at 396 + 11 and c.884 Ma (MSWD ¼ 0.33; Fig. 3c). 86TA100 is K-feldspar augen schist with white mica, quartz,epidote, microcline, biotite and garnet. Four out offive fractions define 404 + 13 and 1149 + 23 Malower and upper intercept ages, respectively(MSWD ¼ 1.9; Fig. 3c). Besides Grenville inheri-tance, these samples indicate Early Devonian mag-matism in the Chortıs Block.

Ar/Ar and Rb–Sr geochronology

We determined ages of deformation and meta-morphic cooling of the Maya and Chortıs Blockrocks by analysis of syn- to post-tectonic minerals.We also dated cooling in pre- to post-tectonicmagmatic rocks, again providing age limits fordeformation.

Rabinal complex. Sample G19s (stop G1, Figs 5a &7a) is weakly migmatitic gneiss from westernGuatemala retrogressed to hornblende–epidoteschist during localized but high-strain, low-T(�400 8C) sinistral shear. Two sericite fractions ofdifferent grain size yielded nearly identical weightedmean ages (WMA). The pooled 67 + 1 Ma age isinterpreted to slightly post-date low-T plasticity inthe shear zones. Sample G27s (stop G57, westernGuatemala, Figs 5a & 7a) is orthogneiss associatedwith amphibolite and shows low-T mylonitizationalong discrete sinistral shear zones; undeformed,probably Miocene (c. 15 Ma, see above) graniteoccurs in the vicinity. The age spectrum is complexand the c. 37 Ma WMA may reflect partially resetLate Cretaceous–Early Cenozoic biotite.

Chuacus complex. Sample 5CO-4b is garnet gneissassociated with K-feldspar–porphyroblastic gneissand schist from the north-central Chuacus complex

Fig. 6. (Continued) (all T– t paths with a merit function value of at least 0.5, Ketcham et al. 2000) are dark grey or black;acceptable-fit solutions (all T– t paths with a merit-function value of at least 0.05) are light grey. (c) AFT age v. elevationplot and elevation–age–distance from the Pacific-coast diagram for a vertical profile and regional samples of thesouth-central Chiapas Massif, southern Mexico. The Tonala shear zone, associated granitoids, metamorphic rocks andmylonites and their ages are indicated (Hbl, hornblende; Bt, biotite); note their influence on one AFT sample. (d) AFTand apatite (U–Th)/He data for two coast-normal sections in southern Mexico from Ducea et al. (2004b) plotted in agev. elevation diagrams; these data are reference sections for our own data and are discussed in the text. Dark grey verticalbars give age range (U–Pb zircon) and intrusion depth of granitoids from about the same area as the thermochronologydata. Light grey bars give interpreted closure through 70–100 8C or c. 4 km depth.


(Fig. 7a); metamorphism evolved through a high-Pstage and amphibolite– to epidote–amphibolite-facies retrogression (see below). Low-Si phengiteand biotite are part of the recrystallized groundmass(�510 8C, c. 0.7 GPa). The phengite spectrum isweakly hump-shaped but the total-fusion (TFA),WMA, and isochron (IA) ages correspond withinerror at 67 + 2 Ma; the biotite is c. 3 Ma younger.A Rb–Sr phengite–plagioclase–whole rock iso-chron is within error of the Ar/Ar ages (Fig. 5b).Epidote–amphibolite-facies retrogression is thusLate Cretaceous. Sample G10s (stop G29, Fig. 7a)is from the southern Chuacus complex and com-prises marble interlayered with garnet micaschistand gneiss just north of El Tambor complex serpen-tinites. The Ar/Ar age of the coarse white mica inthe marble overlaps a Rb–Sr biotite age (G11s;Fig. 7a) of nearby two-mica orthogneiss withinerror at c. 67 Ma (Fig. 5b). Sample G12s (stopG31, Fig. 7a) is fine-grained, mylonitic orthogneiss(c. 1 Ga, unpublished U–Pb zircon) north of G10s.The grain size is due to dynamic recrystallizationof quartz at the transition between subgrain rotationand grain-boundary migration (c. 425 8C, e.g. Stippet al. 2002). White mica is related to the recrystalli-zation fabric, includes albite and is cogenetic with asecond generation of garnet and chlorite; the ubiqui-tous albite may be due to Na-rich fluid exchangebetween the Chuacus complex and the El Tamborcomplex rocks in the hanging wall. Mylonitizationand fluid exchange occurred at or slightly after70 + 3 Ma (Ar/Ar white mica, Fig. 5b). Samples5PB-13b and 5PB-14 typify rocks of the east-centralChuacus complex (Fig. 7b): 5PB-13b is coarse-grained garnet amphibolite associated with garnetmicaschist metamorphosed at c. 540 8C andc. 0.7 GPa (see below). 5PB-14 is hornblende–mica–chlorite schist with quartzitic layers formedat identical PT conditions. Quartz recrystallizeddynamically by subgrain rotation and grain-boundary migration. The hornblende spectrum of5PB-13b is complex with an age component olderthan 137 Ma; several steps are at 65 + 6 Ma. Thelatter are within error of the Ar/Ar (5PB-14) andRb–Sr (5PB-13b) white mica ages (Fig. 5b). Asthe closing temperatures for these systems areclose to the peak-T conditions represented by thesynkinematic minerals, deformation is dated at�65 Ma. Sample 5PB-9b is retrograde biotite–chlorite gneiss from ortho–paragneiss intercala-tions at the eastern end of the Chuacus-complexoutcrop in the Sierra de Las Minas area (below theJuan de Paz serpentinite unit; Fig. 7a). Mylonitiza-tion is upper greenschist facies with basal ,a.slip and grain-boundary migration recrystallizationin quartz and coexisting fracturing and incipientrecrystallization in feldspar. The Ar/Ar white micaspectrum is classic hump-shaped (52–62 Ma)

indicating 40Ar excess, the Rb–Sr white mica–whole rock isochron is at 76.5 + 0.59 Ma (Fig. 5b).Samples 5CO-5a,b, 5CO-6b, 880C, 883B, 885Aand 887C are from the central Chuacus complexand typical for the garnet amphibolite–micaschist–biotite–(?ortho)gneiss–metapegmatite intercalati-ons of this region (Fig. 7a). The pegmatites arevariably deformed but massive stocks are weaklystrained. All rocks share various stages of post-tectonic annealing. The Ar-release spectrum ofphengite of garnet amphibolite 5CO-5a is hump-shaped, indicating 40Ar excess; a low-T plateauis at 225 + 5 Ma (Fig. 5b). Hornblende showsa classic loss-profile spectrum (c. 720–153 Ma).The 70.0 + 1.5 Ma Ar/Ar white mica age fromintercalated quartzofeldspathic paragneiss sample887C expresses Late Cretaceous reheating, alsohighlighted by our U–Pb zircon geochronology(see above). White mica of metapelite sample880C defines a shear zone related to late stagefolding (F2); the individual heating steps cover127–100 Ma. White mica and biotite from para-gneiss sample 883B define a s-c fabric, charac-teristic of late-stage deformation; sinistral flowis constrained at 74 + 2 Ma (white mica) to68.6 + 2.0 Ma (biotite). Sample 854F is migmatiticparagneiss weakly overprinted by late-stage (retro-grade) shear zones from the Rio Hondo area of thesoutheastern Chuacus complex; individual heatingsteps cover 148–94 Ma, probably reflecting amixture of age components. Rb–Sr metapegmatitewhite mica ages are at 76.1 + 1.5 (5CO-5b),57+8 (5CO-6b) and 67 + 11 Ma (885A, extremelycoarse grained mica from the Guatemarmol minewest of Granados), demonstrating Late Cretaceouspegmatite emplacement; our ages confirm thosefrom foliated metapegmatite (K–Ar 73–62 Ma) ofOrtega-Gutierrez et al. (2004). We interpret ourages as Triassic cooling from granite emplacementand associated metamorphism and Late Cretaceousreheating associated with regional metamorphismand local magmatism. The two samples thatsuggest cooling in the Early Cretaceous character-istically are from areas with relative weak late-stagetectonic overprint. On one hand, they (880C)support a pre-Late Cretaceous metamorphism inthe core of the central Chuacus complex, wheresinistral flow is localized and latest Cretaceous(883B). On the other hand, they (854F) demonstratepre-Cretaceous migmatization in the southern partof the eastern Chuacus.

Samples G6s, G7s, 860B and 5C-36a (Figs 5c &7a) are Chuacus-complex rocks immediately belowthe North El Tambor complex rocks. Theselocalities show imbrication of basement (now retro-grade phyllonite, chlorite phyllite, epidotized biotitegneiss, carbonate-bearing garnet–actinolite mica-schist), marble (San Lorenzo marble), serpentinite


Fig. 7. (a) Major geological units of the Motagua suture zone, distribution of Palaeozoic–Cenozoic igneous rocks and new geochronological ages. (b) Definition of mineral cooling and (c) zircon crystallization-age groups in relative probability plots;(d) temperature–time diagrams of all data and (e) selected localities/samples (closing temperatures see text).

and low-grade metasiltstones (Jones Formation).They share a history of extreme fluid flow thatresulted in strong retrogression, local carbonateand albite ‘metasomatism’ and post-tectonicgraben formation, and locally intense quartzveining. White mica of G6s (station G26), retro-gressed basement, has a U-shaped spectrum indicat-ing 40Ar excess with a flat-portion at 60 + 5 Ma(Fig. 5c), comprising most of the released Ar. Retro-grade biotite gneiss G7s (station G27) exhibitscentimetre-sized pull-apart-type tension gasheslocally filled with pseudotachylite that spectacularlydates brittle–ductile, sinistral transtension at c.70 Ma (Fig. 5c). 5C-36a shows coarse white micafrom relict gneiss in overall fine-grained, locallybiotite-rich phyllonite; its 48 + 1 Ma Ar/Ar andthe 35 + 14 Ma Rb–Sr biotite ages are the youngestages encountered in the Chuacus complex rocks andmay reflect long-lasting fluid activity.

Chortıs Block. The following samples characterizemetamorphism and deformation in the Las Ovejascomplex and cooling of associated granitoids(Figs 5d & 7a). Sample 5C-2c is partly myloniticamphibolite intercalated with paragneiss and quart-zite and crosscut by diabase dykes. Hornblende isclearly synkinematic to sinistral shear deformationand shows two high-T steps at .30 Ma and alow-T ‘plateau’ at 20 + 2 Ma; we attribute limitedsignificance to these ages, as only few steps areavailable due to a furnace problem. Sample 5C-5is phyllite faulted against Cenozoic–Quaternaryvolcanic rocks along the Jocotan fault zone. Synki-nematic biotite shows Ar-loss from c. 31 to c. 26 Mawith a mid-T ‘plateau’ at 29 + 2 Ma (Fig. 5d). Thisage indicates Early Oligocene Jocotan fault activity,coeval with mylonitization in the Las Ovejas unit(see below). Samples 5C-23c,d comprise biotitegneiss and amphibolite in a unit of partly migmatiticgneiss, metagabbro and granitic gneiss that istypical Las Ovejas complex lacking retrogression.Synkinematic, fibrous hornblende gives a WMA of35.5 + 1.0 Ma; biotite is at 23.3 + 0.5 Ma andthus .10 Ma younger. Biotite of nearby 5C-33gneiss depicts a loss profile from c. 26 to c. 15 Mawith a mid-T ‘plateau’ at 23 + 2 Ma, identical tobiotite of 5C-23c. K-feldspar shows a loss profilefrom c. 370 to c. 6 Ma; the low-T steps, forming agood-fit IA with atmospheric 40Ar/36Ar at7 + 2 Ma, are difficult to interpret geologically.The high-T steps indicate cooling at 32 + 4 Ma(Fig. 5d). Although from the vicinity of large post-tectonic granite, we interpret our 5C-23 and 5C-33station ages to date amphibolite-facies metamorph-ism and associated migmatization, the minerals lackpost-tectonic annealing and are in part synkinematicto sinistral strike–slip deformation. The amphiboleages closely approximate the age of deformation

(c. 35 Ma); it remains unclear whether the latecooling of these samples through biotite closure(c. 23 Ma) may indicate long-lasting deformation,reactivation, or slow cooling. Samples 5C-26c,dcomprise garnet amphibolite and garnet-bearingbiotite gneiss of the lithological unit describedabove; distinct rocks are migmatitic gneiss andgarnet-bearing leucosome. 5C-26c hornblende isaligned in the foliation (low strain) and exhibitsa loss profile from c. 64 to 17 Ma. Multiple-stepdegassing allowed deconvolution into sub-plateausthat were analysed with isochrons (Fig. 5d).A low-T ‘plateau’ at 18 + 2 Ma has approxi-mately atmospheric 40Ar/36Ar, an intermediate-T‘plateau’ at 21 + 2 Ma is non-atmospheric but exhi-bits a good-fit isochron and a high-T ‘plateau’ at36 + 3 Ma is only weakly non-atmospheric. Theintermediate and low-T ‘plateaus’ are within errorof the 20 + 2 Ma WMA for the mid-T steps of abiotite–minor chlorite mixture; this biotite has aloss profile from c. 28 to c. 3 Ma. The high-T stepsof K-feldspar give 30 + 1 Ma. Sample 5C-37b,again typical Las Ovejas complex of eastern Guate-mala, is amphibolite associated with migmatiticbiotite gneiss and metre-sized bodies of leuco-granite. The hornblende spectrum is complex butcontains a high-T ‘plateau’ with a WMA of39 + 3 Ma, calculated for 40Ar/36Ar of 553, and alow-T ‘plateau’ with a WMA of 18 + 3 Ma, calcu-lated for 40Ar/36Ar of 362; the K–Ca ratios of thelatter suggests a contribution of gas released frombiotite micro-inclusions. 5C-37d biotite revealsAr loss c. 26–9 Ma and a mid-T ‘plateau’ at19 + 2 Ma. The leucogranite gives a white mica–K-feldspar–plagioclase–whole rock Rb–Sr iso-chron of 33.6 + 1.6 Ma. Located in the far-field ofthe large granitoids that were mapped crosscuttingthe Las Ovejas complex, stations 5C-26 and 5C-37typify cooling from amphibolite-facies metamorph-ism at c. 37 Ma and, possibly, relatively slowcooling throughout the Oligocene–Miocene. TheRb–Sr leucogranite white mica age supports thefield observation that these small melt bodies arecogenetic with Late Eocene metamorphism andmigmatization. Sample 5C-28b is pegmatite gneisscrosscutting typical Las Ovejas biotite gneiss; itsRb–Sr whole rock–plagioclase–K-feldspar iso-chron is at 34.2 + 1.0 Ma and a whole-rock–plagioclase–white mica isochron at c. 21 Ma.Sample 5PB-5a is weakly deformed garnet amphi-bolite associated with massive migmatite andsmall orthogneiss bodies and represents a hugexenolith in granite. Peak-PT estimates using core-sections of minerals such as garnet that lack defor-mation are c. 630 8C at 0.7 GPa (see petrology);the analysed hornblende (WMA of 30 + 2 Ma)and biotite (WMA at 28.2 + 0.4 Ma) are synkine-matic matrix minerals that date deformation and


rapid cooling during weak, still amphibolite-faciesretrograde overprint, also detected in the garnetrims; peak-T conditions are at c. 36 Ma (U–Pbzircon, see above).

The next sample group comprises Las Ovejascomplex of northern Honduras and structurally thefootwall block of the Sula graben (Figs 5e & 7a);our dating aimed establishing potential along-strikechanges in deformation/metamorphism/magma-tism and Sula-graben formation. Sample 5H-2a ispost-tectonic, weakly deformed garnet–white mica-bearing pegmatite in migmatitic biotite gneiss andrare leucogranite sills, and 5H-2c is a more stronglydeformed, finer-grained variety of the pegmatite(augen gneiss); these magmatites are outcrop-scaleintrusives. The coarse white mica of 5H-2a is at23.68 + 0.27 Ma (Rb–Sr whole rock–plagio-clase–white mica; the biotite isochron is c. 7 Mayounger) and provides a younger bound for theintrusion of the pegmatite protolith. The whitemica–K-feldspar–plagioclase–whole rock iso-chron of 5H-2c is c. 2 Ma younger (Fig. 5e), prob-ably reflecting the higher strain. Sample 5H-3a issmall, weakly deformed but syntectonic granite instrongly migmatitic biotite gneiss; the 22 + 4 Maisochron age of 10 biotite crystals analysed bysingle-grain laser fusion is .10 Ma younger thanthe intrusion age of the granite (c. 37 Ma, U–Pbzircon, see above). Nearby station 5H-4 typifiesthe heterogeneous lithology of the Las Ovejascomplex in this area. Sample 5H-4b is biotite andhornblende-bearing pegmatite at the rim of asmall, weakly deformed leucogranite body; all mag-matites at this locality are deformed and in severalcases the strain is concentrated in leucocraticdikes. The 38 + 2 Ma hornblende is identical tothe age of the concordant zircon fraction of nearbygranite gneiss 92RN412 (see above); biotite is.10 Ma younger at 25 + 3 Ma (laser single-grainages). Sample 5H-4c shows synkinematic horn-blende and biotite from a sinistral transtensive,greenschist-facies shear zone in amphibolite: horn-blende is at 30 + 2 Ma, biotite show a 25 + 2 Mamid-T ‘plateau’ (Fig. 5e); these ages best date for-mation of the Sula graben (see also below). Largehornblende crystals of sample 5H-4d, the weaklydeformed host amphibolite to the above shearzone, contain small hornblende crystals and arewithin error (29 + 3 Ma) of the synkinematicsample 5H-4c amphiboles. Sample 5H-9a, ortho-gneiss and migmatite that yielded Permo-TriassicU–Pb zircon ages (see above) and are distinctlydifferent from the Las Ovejas complex rocks,gives Rb–Sr whole rock–plagioclase–white micaand biotite isochrons of c. 162 and c. 135 Ma, res-pectively. These dates have two possible inter-pretations: either they reflect very slow coolingfrom the Permo-Triassic metamorphic/magmatic

event or Middle Jurassic reheating not reflected inthe zircon geochronology in this sample butdetected in the Las Ovejas complex of Guatemalaand in the central Chortıs Block of Honduras.We prefer the latter interpretation, as we considerextremely slow cooling unlikely in the active tec-tonic setting of Early Mesozoic northern CentralAmerica.

Sanarate complex. Samples G3s and G5s (stops G23and G24; Figs 5f & 7a) are phyllite and quartz-bearing amphibolite from the Sanarate complex,respectively. Quartz shows low-T plasticity withgrain-growth inhibition and some post-tectonicannealing; together with hornblende and whitemica it is synkinematic in discrete mylonite zones.The white mica spectrum of G3s shows Ar-lossfrom c. 195 to c. 137 Ma with a mid-T ‘plateau’ at155 + 10 Ma, and G5s hornblende from c. 157 toc. 40 Ma; a high-T ‘plateau’ is at 156 + 5 Ma andlow-T steps at �40 Ma. We interpret these data asindicating Middle–Late Jurassic deformation andCenozoic reactivation.

Granitoids of the Chortıs Block. We dated coolingin a few of the mostly undeformed granitoids ofthe northern Chortıs Block (Figs 5g & 7a). Sample5C-11, granodiorite of the Chiquimula batholith,shows biotite partly altered to chlorite; 10 grainsdated by laser fusion give a 90 + 10 Ma WMA.Eight out of 12 K-feldspar grains of sample5C-21, sub-volcanic granite also from the Chiqui-mula batholith, have a 20 + 1 Ma laser-fusionWMA; this demonstrates a composite nature ofthis batholith. K-feldspar of sample 5H-5, granitefrom the rift flank of the Sula graben, shows Arloss from c. 37 to c. 28 Ma; the low-T ‘plateau’ at28.3 + 0.5 Ma indicates rapid cooling (Fig. 5g). Inline with the synkinematic hornblende and biotiteages reported from this area (see above), this agesuggests initiation of formation of the Sula grabenat 28 + 3 Ma. Six biotite grains of sample 5H-13,granite from south of the Las Ovejas unit, gave alaser-fusion age of 59 + 3 Ma. The lack ofmodern U–Pb geochronology still impedes clearage assignments of plutonic events in the ChortısBlock. Because of the often sub-volcanic nature ofthese granitoids, we assume that the available geo-chronological data reflect, to first-order, theseevents. Figure 5h pools all granitoid data andgives first-order groupings: there is significant mag-matism at c. 36 Ma starting �40 Ma, at c. 59 Ma,and less significant at c. 81, c. 120, c. 161 andc. 260 Ma.

Southern Mexico. For correlative purposes (seediscussion), we analysed samples from southernMexico with the Ar/Ar method (Figs 5i & 1c).


Sample ML-18 is Magdalena migmatite of theeastern Acatlan complex (Fig. 2g). Biotite showsAr loss from c. 163 to c. 88 Ma with a high-T‘plateau’ at 162 + 2 Ma; the latter traces coolingfrom migmatization and San Miguel dyke intrusion(175–171 Ma U–Pb ages; e.g. Keppie et al. 2004;Talavera-Mendoza et al. 2005). White mica fromsample ML-37, mylonitic gneiss of the southern-most Oaxaca complex (Fig. 1c), shows Ar lossfrom a high-T ‘plateau’ at 315 + 3 Ma to a low-T‘plateau’ at 247 + 3 Ma; the latter corresponds toages of Permo-Triassic arc magmatism/meta-morphism, exemplified by the Chiapas Massif ofsoutheastern Mexico (e.g. Weber et al. 2006), theformer corresponds to a widespread tectonothermalevent in the Acatlan complex of southern Mexico(e.g. Vega-Granillo et al. 2007). Biotite fromsample MU-12, migmatitic gneiss at Cruz Grande(Xolapa complex, Fig. 1c) yielded 29 + 1 Ma(WMA), in line with cooling from widespread c.35–30 Ma Xolapa plutonism (Herrmann et al.1994; Ducea et al. 2004a). Sample MU-13 is LasPinas mylonitic orthogneiss, sampled east of LaPalma, which forms the structural top of theXolapa complex. Solari et al. (2007) reported a54.2 + 5.8 Ma intrusion age (U–Pb zircon), and50.5 + 1.2 and 45.3 + 1.9 Ma cooling ages (K–Ar and Rb–Sr biotite, respectively) from a moder-ately deformed variety. Our sample was specificallytaken to date low-T, high-strain mylonitic flow (c.300 8C) along the Tierra Colorada (Riller et al.1992) or La Venta (Solari et al. 2007) top-to-NW,sinistral-transtensional shear zone; synkinematicbiotite is at 35 + 1 Ma. Sample ML-39, myloniticgneiss at Pochutla, is from a narrow, late-stageshear zone in the regional Chacalapa shear zonealong the northern edge of the eastern Xolapacomplex; its major sinistral strike–slip stage is con-strained at 29–24 Ma (Tolson 2005; Nieto-Samaniego et al. 2006). Our 15 + 2 Ma biotiteage suggests reactivation, also indicated by faultingin the 29 + 1 Ma, post-mylonitic Pochutla granite(Herrmann et al. 1994; Meschede et al. 1996).

Samples PA3-4, PI28-1, CB31, CB35 and CA34are from the Chiapas Massif of southeastern Mexico(Figs 1c & 5i). A belt of small plutons along thesouthwestern margin of the massif is associatedwith metamorphic rocks that are partly mylonitizedand faulted along a NW-trending, .100 km longzone, named Tonala shear zone by Wawrzyniecet al. (2005). The late ductile and brittle kinematicsis sinistral strike–slip, overprinting an older,higher-T fabric; the zone probably connects tosimilar sinistral strike–slip fabrics and plutonsalong the Polochic fault (Fig. 1c; see below). Ourgoal was to obtain ages for the intrusions andsinistral shear. Hornblende of diorite CB31 andtonalite CA34 yielded 15.5 + 1.5 and 11 + 2 Ma,

respectively; the former and latter are similar to14.65 + 0.42 granite GM14c along the Polochicfault zone of westernmost Guatemala (see above)and a 10.3 + 0.3 Ma, pervasively sheared pluton(U–Pb zircon, Wawrzyniec et al. 2005) of theTonala shear zone, respectively. Hornblende andbiotite from three mylonitic gneisses (PA3-4,CA35, PI28-1) gave within error identical agesat c. 8.2 Ma, interpreted to date sinistral sheardeformation. One hornblende age (CB35) at28.5 + 1.5 Ma is speculatively interpreted asapproximating the age of the higher-T flow/meta-morphism, in line with Oligocene deformationages along the northern margin of the Xolapacomplex (see above).

Fission-track thermochronology

Our TFT and AFT thermochronology is only recon-naissance in nature (Figs 6 & 7a). The two TFT agessupport the available Ar/Ar and Rb–Sr geochronol-ogy, being close to biotite cooling ages. The AFTages from Guatemala and Honduras show positiveage v. elevation correlations both in the centralChuacus complex and the central and eastern partsof the northern Chortıs Block (Fig. 6a); exhumationrates, if significant, are similar and slow for the areasnorth and south of the Motagua fault zone (c. 0.038and 0.035 mm/annum, respectively), and possiblyslightly more rapid along the western Sula-grabenflank of northern Honduras (c. 0.046 mm/annum).The Chuacus-complex ages depict a cluster atc. 30 Ma, the low elevation ages from the LasOvejas complex at c. 12 Ma (Fig. 6a). Despitehigh elevation, the youngest ages north of theMotagua fault zone occur in western Guatemalaalong the Polochic fault in a small, c. 15 Magranite and the mylonitic augengneiss it intrudes(samples GM13s, GM14s, Fig. 7a). The c. 2 Maage difference between crystallization and coolingthrough c. 100 8C of the granite reflects quenchingat high crustal levels. The c. 11.5 Ma TFT and thec. 4.8 Ma AFT ages from low-T mylonite approxi-mate the age of ductile–brittle to brittle Polochicfaulting. The outcrop conditions prevent assessmentof whether the apparently undeformed granite islocally affected by mylonitization but both myloniteand granite show east-trending, sinistral faults. Theoldest ages at high elevation in the Chortıs Blockcluster at c. 22.5 Ma. The trend in the northernChortıs Block for AFT ages to become youngertoward major fault zones (Motagua fault, Sulagraben), with ages as young as 7.8–3.8 Ma,(Figs 6a & 7a), may be apparent, as these samplesalso occupy the lowest elevations and thus are cer-tainly also an expression of the age v. elevationtrend. The presence of these ages at low elevationsalong the Motagua fault suggests that its current


activity (e.g. M7.5 1976 Guatemala earthquake,incorporating strike–slip and normal components;Plafker 1976) extended at least into the latestMiocene–Pliocene.

The track-length distributions of our samples areunimodal, narrow, and symmetric. Modelled T– tpaths give monotonous continuous-cooling typesolutions (Fig. 6b). All T– t models (except thec. 15 Ma granite from western Guatemala, seeabove) show a phase of rapid cooling in the lastfew Ma that is an artefact caused by annealing atambient temperatures acting over geological time.Low-T track-length reduction has been describedfor fossil tracks in age standards (e.g. Donelicket al. 1990) and borehole samples (Jonckheere &Wagner 2000). This reduction is not incorporatedinto the annealing equations derived from labora-tory experiments on induced fission tracks, whichaccount for annealing processes that take placewithin the partial annealing zone (Jonckheere2003a, b).

Our samples from the Chiapas Massif, southeast-ern Mexico, comprise a �1.5 km elevation profile(Fig. 6c) and two swaths separated by ,40 kmwith little elevation difference; no apparent majorfault separates these samples. The youngestsample is at c. 15.7 Ma, close to the Tonala shearzone and the c. 11 Ma tonalite (Fig. 6c); it is mostlikely thermally influenced, supporting theMiocene magmatism and deformation along thisshear zone. The remaining samples range fromc. 25 to c. 39 Ma without a clear age v. elevationtrend and a weighted average age of 30.4 +2.5 Ma; we suggest that these samples indicate rela-tively rapid cooling at c. 30 Ma.

Interpretation of new geochronology

Our U–Pb zircon geochronology outlines EarlyPalaeozoic to Late Cenozoic high-T metamorphism,migmatization, and magmatism (Figs 3 & 7c). In theRabinal complex of the Maya Block, we resolvedprolonged Ordovician–Silurian (c. 410–485 Ma,age groups at 413 and 465 Ma), Permo-Triassic(c. 215–270 Ma, main age group at 218 Ma), andMiddle Jurassic (c. 163–177 Ma, main age groupat 171 Ma) events (Fig. 3a); the Triassic eventincludes migmatites. Miocene (c. 15 Ma) graniteintruded into shallow crustal levels (nearly identicalU–Pb zircon and AFT ages) in western Guatemala.Chuacus-complex rocks inherited Devonian zircons(c. 403 Ma) and experienced Triassic (c. 239–211 Ma, age groups at 238 and 218 Ma) magmatismand Cretaceous high-grade metamorphism (c. 70–78 Ma, main age group at 74 Ma; Th–U , 0.03).These events are partly supported by Ar/Ar andRb–Sr cooling ages (c. 238–212; c. 161; c. 75–65 Ma). The Late Cretaceous metamorphic event

is about coeval with pegmatite emplacement atc. 74 Ma (Rb–Sr and K–Ar on very coarse-grainedwhite-mica). Local Early Cretaceous cooling ages(two samples) are probably geologically meaning-less, possibly reflecting pre-Cretaceous eventsoverprinted by Late Cretaceous deformation. Mig-matization in the Rıo Hondo area is pre-Cretaceousand it remains to be demonstrated that it is Triassic,as in the Rabinal complex of western Guatemala;lithology, structure (see below) and our Ar/Ardata suggest that this region may constitute adistinct unit.

In the Las Ovejas complex of the northernChortıs Block and in the central Chortıs Block, weresolved Early Permian (c. 264–283 Ma, main agegroup at 273 Ma) magmatism followed by EarlyTriassic (c. 242–253 Ma, main age group at245 Ma) high-grade metamorphism (including mig-matite) that is also reflected by a K–Ar biotitecooling age (222 + 8 Ma; MMAJ 1980), andMiddle to Late Jurassic (c. 140–191 Ma, agegroups at 159 and 166 Ma) magmatism and sub-sequent cooling (Rb–Sr and Ar/Ar ages at c.155 Ma). Overwhelmingly, the northern margin ofthe Las Ovejas complex is dominated by LateEocene high-grade metamorphism, migmatization,and plutonism at 40–35 Ma (age group at 37 Ma).Cooling ages (Rb–Sr, Ar/Ar) related to this per-vasive event cover c. 35–20 Ma and also comefrom Roatan Island (c. 36 Ma). Our geochronologyalso supports Silurian (c. 400 Ma) and Jurassic(c. 167 Ma) magmatism, and likely a Permo-Triassicevent both in the northern and central Chortıs Block.Based on U–Pb zircon geochronology, the Mayaand Chortıs Blocks share Ordovician to Jurassicevents (Fig. 7c; the evident Proterozoic similarityis not addressed here). The northern and centralChortıs Blocks are distinctly different from theRabinal and Chuacus complexes in the Late Cretac-eous and Early Cenozoic: Late Cretaceous high-grade metamorphism only occurs in the Chuacuscomplex and Eocene magmatism and metamorph-ism only can be found in the northern Chortıs Block.

The most striking, first-order result showcasedby our Ar/Ar and Rb–Sr geochronology is theclear-cut separation of the Rabinal and Chuacuscomplexes from northern Chortıs Block reflectedin all cooling ages (Fig. 7a & b): the Rabinal andChuacus complexes, rimmed by the Polochic andthe Motagua fault zones, cooled through c. 500 8Cand through 275–350 8C c. 40 Ma earlier than theLas Ovejas complex south of the Motagua faultzone; cooling through c. 100 8C occurred c. 15 Maearlier in the north than the south. Obviously, therock units north and south of the Motagua faultzone experienced distinctly different thermalevolutions during the Late Cretaceous and Ceno-zoic: the Chuacus complex was heated to high


amphibolite-facies metamorphism (see below) andcooled relatively rapidly (c. 28 8C/Ma) to uppercrustal temperatures during the Late Cretaceous(Fig. 7d); subsequently, cooling (c. 6 8C/Ma) andexhumation was apparently slow (c. 0.038 mm/annum; see Figs 6a & 7d). Multi-mineral T– tpaths from central and east-central Chuacus sam-ples (Fig. 7e) support this two-stage history. TheLas Ovejas complex lacks Cretaceous regionalmetamorphism; its cooling from �40 Ma appearssteady-state at c. 16 8C/Ma (Fig. 7d & e). Our geo-chronology does not prove whether slow coolingactually occurred, thus potentially reflecting long-term wrenching, or represents deformation andcooling at 40–35 Ma and reactivation at 25–18 Ma. As we found clear synkinematic mineralgrowth defining both age groups (see below), wefavour the latter scenario. Along the southernmargin of the Chiapas Massif, U–Pb and Ar/Argeochronology demonstrates Cenozoic plutonism(c. 29, c. 16, c. 11 Ma) and sinistral strike–slip atc. 8 Ma along the Tonala shear zone.

Most AFT ages within the Rabinal and Chuacuscomplexes group around c. 30 Ma and thus corre-spond to the c. 30 Ma age cluster in the ChiapasMassif. We suggest that these ages indicate athermal disturbance at �30 Ma induced by Ceno-zoic arc magmatism and sinistral displacementbetween the Maya and Chortıs Blocks alongand across this arc (see below). The young TFTand AFT ages along the Polochic fault zone relateits sinistral strike–slip deformation to the LateMiocene–Recent deformation along the Tonalashear zone, thus supporting their contemporaneousactivity (Fig. 1c). These ages also support Burkart’s(1983) estimate of major displacement (c. 130 km)along the Polochic fault between 10 and 3 Ma.The presence of very young AFT ages along theMotagua fault zone extends its neotectonic activityinto the latest Miocene–Pliocene. The cluster ofAFT ages at c. 22.5 Ma in the northern ChortısBlock probably reflects continuation of the coolingfrom Late Eocene–Early Oligocene magmatismand metamorphism; the c. 12 Ma cluster is attribu-ted to a thermal event that is connected with theflare-up of magmatism (ignimbrite event) alongCentral America (see below).

Taking the peak-P conditions of c. 0.7 GPa (seepetrology below) reached in the Chuacus and LasOvejas complexes during their Cretaceous(Chuacus complex, c. 75 Ma) and Cenozoic (LasOvejas complex, c. 40 Ma) metamorphism, itappears that (i) early average exhumation rateswere at least one order of magnitude more rapidthan those derived from the AFT age v. elevationrelationship (Fig. 6a); (ii) the presence of AFTages at high elevations, which approach those ofthe Ar/Ar and Rb–Sr biotite geochronometers,

suggests that the change from relatively rapid torelatively slow rates occurred relatively early inthe exhumation process (c. 60 Ma in the Chuacuscomplex; c. 20 Ma in the Las Ovejas complex).The fact that the late-stage exhumation rates (AFTage v. elevation data) are grossly similar in theChuacus and Las Ovejas complexes suggests,given the strongly different cooling rates, muchhotter crust in the Las Ovejas complex.

The difference between the southern Maya (hereincluding the Chuacus complex) and northernChortıs Blocks during the Cretaceous–Cenozoic isemphasized by the distinctly different P–T–d– tconditions encountered in the North and South ElTambor complex allochtons (Harlow et al. 2004;Tsujimori et al. 2004): extremely low-T lawsoniteeclogites experienced blueschist-facies retrogres-sion at c. 120 Ma south of the Motagua fault,whereas higher-T zoisite eclogites show blueschist-facies retrogression at c. 76 Ma north of the fault.These P–T–d– t conditions illustrate differencesin tectonic evolution and time of the emplacementof these subduction–accretion complexes. Signifi-cantly however, in the high-P/low-T case, theregion south of the Motagua fault zone cooledearlier and cooling in the North El Tamborcomplex is coeval with the onset of cooling in theChuacus complex.

New petrology

We selected 12 representative thin sections of mag-matic and metamorphic rocks from both sides of theMSZ for electron microprobe analysis. Mineralabbrevations taken from Kretz (1983). Mineralassemblages and summary of the P–T– t estimatescan be found in the Supplementry materials, andFigure 8 shows the PT space ofour and literature data.

Rabinal complex, western Guatemala

Strongly foliated hornblende–epidote schist G19scontains phengite (Si-content c. 3.3 p.f.u.), greenamphibole (edenite to ferro-edenite associatedwith magnesio- to ferro-hornblende), epidote(XPs � 0.15; cf. Dahl & Friberg 1980), and saus-suritized plagioclase with an An-content of0–2 mole%. Biotite forms post-tectonic, topotacticflakes on phengite. Accessory minerals are opaquesand titanite. Phengite geobarometry yielded.0.7 GPa at 500 8C (Massonne & Szpurka 1997).The calibrations of Spear (1981) and Plyusnina(1982) gave 0.7–0.8 GPa at 450–500 8C, whichare reproduced by Colombi’s (1988) Ti-in-hornblende geothermometer. Our 67 + 1 MaAr/Ar phengite age demonstrates latest Cretaceousexhumation through the middle crust with anaverage rate c. 0.37 mm/annum. Epidote schist


G21s is mylonitic with numerous porphyroclastsof epidote (XPs core ¼ 0.295, XPs rim ¼ 0.215),albite (An0.3 – 0.9) and microcline (Kfs98Ab1 – 2

Cel1 – 2An,0.5). The feldspars lack recrystallizationbut quartz recrystallized dynamically by subgrainrotation. Greenish biotite (XMg c. 0.47) and

phengite (c. 3.3 Si p.f.u.) are common; opaques,titanite, apatite and zircon are accessories.Strong mineral chemical disequilibria betweenthe phases preclude PT estimates; the Si-contentof phengite suggests c. 0.8 GPa (Massonne &Szpurka 1997).

Fig. 8. Pressure–Temperature–time diagrams: (a) Pressure–temperature–time space of our new petrology andgeochronology, and literature data of Ortega-Gutierrez et al. (2004). A few Chuacus-complex samples plot in theeclogite field. They follow an isothermal decompression path to amphibolite facies. The eclogite-facies metamorphismis at best constrained to between c. 440 and 302 Ma (see text). Most of the complex passed through epidote–amphibolitefacies metamorphism and subsequent retrogression to greenschist facies; these events are well-dated as Late Cretaceous(�75 Ma). The Sanarate complex of the northwestern Chortıs complex gives Jurassic (c. 155 Ma) epidote–amphibolite-facies metamorphism that evolved prograde out of greenschist facies. The Las Ovejas complex (ChortısBlock) shows Cenozoic (�40 Ma) epidote–amphibolite to amphibolite facies. (b) Motagua suture zone pressure–temperature–time evolution from this study and Tsujimori et al. (2004, 2005, 2006), Harlow (1994), Harlow et al.(2004) and Brueckner et al. (2005). The Late Cretaceous Chuacus complex and Mid-Cenozoic Las Ovejas complexpressure–temperature fields are from this study. The northern zoisite eclogites (on Maya Block) formed at c. 131 Maand reached mid-crustal conditions at about the same time as the Chuacus rocks experienced reheating at epidote–amphibolite-facies conditions accompanied by local magmatism. We suggest that the two are linked to the sameoceanic-to-continental subduction zone (see text). The southern lawsonite eclogites (on Chortıs Block) formed atc. 132 Ma and were at mid-crustal conditions already by c. 120 Ma. Their emplacement is probably not geneticallylinked to the high-grade metamorphism and magmatism of the Las Ovejas rocks, which is a result of theCaribbean Plate translation and the formation of the Cayman Trough (see text).


Fig. 9. Overview map (top left) locates geological–structural map (bottom) in the Maya Block of western Guatemala. The latter compiles and reinterprets available 1 : 50 000 geological maps and plots the main structural features both on map and in stereonets(lower hemisphere, equal area); it also serves as legend for the following structural data maps. LPO (lattice preferred orientation) of quartz c-axis; D, deformation event; X, Y, Z, principal axes of finite strain. Fault-slip data and principalstress orientations 1–3: faults are drawn as great circles and striae are drawn as arrows pointing in the direction of displacement of the hanging wall. Confidence levels of slip-sense determination are expressed in the arrowhead style: solid, certain;open, reliable; half, unreliable; without head, poor. Arrows around the plots give calculated local orientation of subhorizontal principal compression and tension.

Chuacus complex, central Guatemala

Garnet gneiss 5CO-4b is dominated by orthoclase(Kfs93Ab5Cel1An,0.5) and minor, partly saussuri-tized plagioclase (An0.2 – 1.4). Quartz showssubgrain-rotation recrystallization. Chloritizedbiotite (XMg ¼ 0.17) and phengite have epidote(XPs ¼ 0.23) and plagioclase inclusions. Phengitehas elevated Si-contents of up to 3.50 p.f.u. in thecore, indicating high-P conditions, whereasSi-contents as low as 3.27 p.f.u. occur in the rim sec-tions. Garnet forms porphyroblasts with inclusionsof plagioclase, quartz, minor ferro-hornblende/ferro-tschermakite, omphacite (10 mole% jadeite)and rutile. It displays compositional zoning with abell-shaped spessartine curve (core Alm43GAU44-Sps10Prp0.5; rim Alm47GAU48Sps5Prp0.5). Carbon-ate, zircon and titanite form accessory minerals inthe matrix. Relict phases highlight the polymeta-morphic evolution of 5CO-4b. Assuming equili-brium between the omphacite inclusion in garnet,the adjacent garnet core section and the core ofthe matrix phengite, PT conditions of c. 583 8Cand c. 2.04 GPa are calculated using the garnet–phengite–omphacite geobarometer of Waters &Martin (1993) (phengite activity model of Holland& Powell 1998), clinopyroxene activity model ofHolland (1980, 1990), and garnet activity model ofGanguly et al. (1996) in combination with thegarnet–clinopyroxene geothermometer of Krogh(2000) (calculation of ferric iron in clinopyroxeneaccording to Ryburn et al. 1976). Similar tempera-tures are calculated for high-P phengite and garnetcore analyses using the thermometer of Green &Hellman (1982). These PT conditions, however,only indicate a point on the prograde PT path as thehighest pressure is indicated by the highest aGrs

2*aPrp

(Schmid et al. 2000) at an intermediate positionbetween garnet core and rim. The strong retrogradeoverprint is mirrored by pervasive growth of matrixbiotite and low-Si phengite. Outermost garnet rims(highest XPrp) and adjacent biotite yielded c. 510 8C(geothermometer of Hodges & Spear 1982) atc. 0.7 GPa (barometry of Massonne & Szpurka1997) using paragenetic phengite; this demonstratesre-equilibration under epidote–amphibolite-faciesconditions. Because of this overprint, feldspar ther-mometry (Fuhrman & Lindsley 1988) on coexistingplagioclase (An0.2 – 1.4) and K-feldspar (integratedanalysis with a 50 mm beam diameter) indicatescooling temperatures of 400–440 8C. The lowerintercept U–Pb zircon age indicates that high-Pmetamorphism is ,638–477 Ma and our Ar/Arand Rb–Sr phengite and biotite ages date the latestportion of the retrograde path at 68–64 Ma.

Garnet amphibolite (eclogite) 5CO-5a containsminor quartz showing subgrain-rotation recrystalli-zation and foliation-defining, large calcic amphibole

(pargasite to edenite). Omphacite with 35 mole%jadeite at the rim to 40 mole% in the core ispresent as relicts in garnet and hornblende coresand less commonly in the matrix. Well-alignedphengite is the dominant sheet silicate; Si-contentdecreases from core (3.32 p.f.u.) to rim (3.27p.f.u.). Garnet displays weak zoning with decreasinggrossular and pyrope contents and increasingalmandine contents from core (Alm60.5GAU23-Sps0.5Prp16) to rim (Alm63GAU20.5Sps0.5Prp15.5).Plagioclase (An2 – 4) is secondary, forming alongthe rims of garnet and hornblende; irregular spotsof An10 – 14 are also present. Late-stage biotite(XMg c. 0.53) grew at the expense of phengite.Apatite, rutile, ilmenite and pyrrhotite are acces-sories. PT conditions of the eclogite facies eventwere calculated applying the garnet–phengite–omphacite geobarometer of Waters & Martin(1993) (activity models see sample 5CO-4b) incombination with the garnet–clinopyroxenegeothermometer of Krogh (2000). Using themethod of Ryburn et al. (1976) to estimate Fe3þ

in clinopyroxene, PT conditions of c. 675 8C atc. 2.4 GPa were calculated using the core compo-sition of garnet (highest aGrs

2*aPrp), the omphaciteinclusion in garnet with the highest jadeitecontent, and the phengite with the highest Sicontent (Carswell et al. 2000). The application ofother methods to estimate the ferric iron content inomphacite (e.g. Carswell et al. 2000; Schmid et al.2000; Krogh 2000; Krogh & Terry 2004) results inan uncertainty of +50 8C for the geothermometry.Omphacite in the matrix probably was affected byretrograde processes and consequently shows awide range of temperatures (480–690 8C). Temp-eratures for retrograde overprint of the eclogitewere estimated at 630–690 8C using the garnet–hornblende thermometer of Graham & Powell(1986) for garnet rims and adjacent hornblende.The GRIPS geobarometer of Bohlen & Liotta(1986) calculates pressures of c. 1.1 GPa for theseminerals and associated plagioclase, rutile andilmenite. Slightly higher pressures of 1.2–1.4 GPaare calculated using the garnet–amphibole–plagioclase geobarometer of Kohn & Spear (1990).These estimates indicate isothermal decompressionduring exhumation of the eclogite. Ortega-Gutierrezet al.’s (2004) c. 302 Ma lower intercept zircon agefrom a leucosome interbedded with relict eclogitefrom the area of 5CO-5 was interpreted to providea minimum age for the Chuacus high-P event.

The strong foliation of garnet-bearing micaschist5PB-13a is formed by alternating mica- and quartz-rich layers; quartz shows subgrain-rotation andgrain-boundary-migration recrystallization. Small,almandine-rich and spessartine-poor garnet grains(Alm70GAU10Sps1Prp18) are included in largealbite porphyroblasts (An0.2 – 0.7), and larger garnet


porphyroblasts with Alm74GAU2Sps13Prp8 crystal-lized within the matrix; the garnets are slightlyinhomogeneous without regular zonation patterns.Green to brownish biotite (XMg c. 0.58) is partlyaltered to oxychlorite. Si-content in phengitedecreases from core (3.3 p.f.u.) to rim (,3.2p.f.u.). This correlates with decreasing Ti-contentand indicates a retrograde evolution. Because ofthe lack of critical mineral assemblages, only afew thermobarometers can be used to evaluate thePT conditions of these schists. The Si-content ofphengite yields �0.6–0.8 GPa; the garnet–biotitegeothermometer (Hodges & Spear 1982) and thegarnet–phengite geothermometer (Green &Hellman 1982) point to 560–580 8C for the garnetinclusions in albite and 460–475 8C for the matrixgarnets. Garnet amphibolite 5PB-13b is well foli-ated, medium- to coarse-grained, and compriseslarge, pale-green calcic amphibole (magnesio-hornblende to edenite) with numerous garnet andquartz inclusions; the matrix has also phengite andclinozoisite. Phengite Si-content (3.3–3.15 p.f.u.)and XMg (0.76–0.66) decrease toward the rims.Clinozoisite (XPs ¼ 0.08–0.09) forms largeprisms. Quartz shows both dynamic subgrain-rotation and grain-boundary-migration recrystalli-zation. Albite (An2 – 6) has numerous quartzinclusions. Small, euhedral garnet yields Alm55-GAU35Sps5Prp4 in the core and Alm54GAU30Sps5-Prp11 in the rim and has increasing XMg valuesfrom core to the rim, indicative of progradegrowth. One garnet differs compositionally, withhigh XMg in the core and abruptly increasing alman-dine content at the expense of grossular and spessar-tine towards the rim; this may indicate an earlier,prograde metamorphic evolution. Retrograde chlor-ite is widespread; apatite, zircon, and titanite areaccessories. Pressures of �0.7 GPa (Massonne &Szpurka 1997) at 530–550 8C (garnet–hornblendegeothermometer of Graham & Powell 1986) areobtained; the Ti-in-hornblende thermometer ofColombi (1988) provides c. 590 8C. Our Ar/Aramphibole and Rb–Sr phengite ages suggest meta-morphism at c. 65 Ma at this locality.

Hornblende–mica–chlorite schist 5PB-14shows distinct layering with amphibole–mica andquartz that is dynamically recrystallized by subgrainrotation and grain-boundary migration. Pseudo-morphs of chlorite (ripidolite after Hey 1954) afterrelict garnet and lamellae of oligoclase (An12)within albite (An0.5 – 2) demonstrate retrogradeoverprint. Green amphibole (tschermakite to mag-nesiohastingsite) and clinozoisite (XPs c. 0.13) aresubordinate. White mica shows increasingSi-content from the core to the rim (3.12–3.20p.f.u.). Pale brown biotite is rare. Carbonate isFe-bearing dolomite. Opaque minerals, rutile andtitanite form accessory phases. PT conditions of

500–550 8C at c. 0.8 GPa are suggested by thegeothermobarometer of Plyusnina (1982). This esti-mate is reproduced by the phengite geobarometry(Massonne & Szpurka 1997) and the amphibole–plagioclase geothermometers of Spear (1981),Blundy & Holland (1990), and Holland & Blundy(1994). A temperature of 600–630 8C is indicatedby the Ti-in-hornblende geothermometer ofColombi (1988), probably due to the high bulkTi-content of the rock. This may, however, alsoindicate an early stage of medium- to high-grade,amphibolite-facies metamorphism. Our 65 + 1 MaAr/Ar white mica age dates cooling duringgreenschist-facies retrogression.

Gneiss G14s is inhomogeneous with alternatingplagioclase–quartz and mica–hornblende layers.Plagioclase is albite (An0.5 – 2); epidote (XPs ¼ 0.2)and chlorite are present in the mafic layers. Phengitehas Si-contents of 3.27–3.34 p.f.u. Biotite yieldsXMg of c. 0.51. Green amphibole is magnesio- andferro-hornblende. Small garnet within the maficlayers and as tiny inclusions in plagioclase hasa bell-shaped spessartine curve and rimwarddecreasing XFe and XMn; cores and rims are Alm40-GAU31Sps25Prp1 and Alm48GAU37Sps13Prp2, res-pectively. Accessories are carbonate, opaques andtitanite. Triboulet’s (1992) geothermobarometercalculates c. 610 8C at c. 0.62 GPa for the rim com-position of the minerals, supported by phengite geo-barometry at 0.6–0.7 GPa at 500–600 8C. About550 8C is obtained by the garnet–hornblende(Graham & Powell 1986) and the garnet–chloritegeothermometer (Ghent et al. 1987). These PTestimates indicate peak conditions at the transitionfrom epidote-amphibolite to amphibolite-faciesmetamorphism.

Intrusive sequence, Juan de Paz ophiolite,

North El Tambor complex

Fine- to medium-grained microdiorite 5PB-8 showsagglomeration of hornblende and plagioclase.Plagioclase shows normal zoning with An-contentdecreasing from 85 to 55 mol% from core to rim.Calcic amphibole (magnesiohornblende) is slightlybrownish and short prismatic. A few clinopyroxenerelicts yield 95–98 mol% augite (end membercalculation after Banno 1959) and are mantledby hornblende. Opaque minerals and apatiteare common accessories. Anderson & Smith’s(1995) Al-in-hornblende geobarometer indicatesc. 140 MPa at 730–860 8C, calculated with theTi-in-hornblende and the plagioclase–hornblendegeothermometers. Bertrand et al. (1978) dated alikely metamorphic overprint in related doleritesfrom the Baja Verapaz (Fig. 1c) at c. 75 Ma (seeabove).


Sanarate complex

Chlorite–epidote–amphibole schist 5G-1a containsalbite (An1 – 3.5), clinozoisite (XPs � 0.1), chlorite(ripidolite), amphibole (edenite) and phengite(3.2–3.3 Si p.f.u.). The application of the geother-mobarometer of Triboulet (1992) yields peak PTconditions of c. 550 8C at c. 0.65 GPa for the rimsof the minerals. The cores indicate c. 380 8C atc. 0.55 GPa, outlining a prograde metamorphicloop. These data highlight the transition fromlower greenschist to epidote–amphibolite faciesand are supported by the geothermometry ofColombi (1988) and Holland & Blundy (1994),which provide c. 415 8C for the core and c. 560 8Cfor the rim of the amphibole. Si-contents of syntec-tonic phengite accordingly yield 0.6–0.65 GPa(Massonne & Szpurka 1997). A likely age of thismetamorphism is Middle Jurassic, given by thehornblende and white mica ages from adjacentlocations G23, G24, which comprise similar rocks.

Las Ovejas complex

Garnet amphibolite 5PB-5a displays a fine- tomedium-grained matrix with synkinematic biotite,amphibole and large, poikiloblastic garnet. Matrixquartz shows numerous subgrains. Biotite is Ti-rich(up to 4.1 wt% TiO2) and has an XMg of about 0.24.The green amphibole is hastingsite. Large plagio-clase grains are chemically homogeneous withAn-content of c. 35 mol%. Garnet cores showAlm69GAU19Sps8Prp4; XFe and XMn increasetowards the rims, indicating retrograde overprint.Accessories are opaques and apatite. Garnet–hornblende geothermometry (Graham & Powell1986) in combination with garnet–amphibole–plagioclase–quartz geobarometry (Kohn & Spear1990) yields 627–655 8C at 0.7–0.8 GPa. Com-parable temperatures are calculated with garnet–biotite thermometers of Hodges & Spear (1982),Indares & Martignole (1985) and Perchuk & Lavren-t’eva (1983). Ar/Ar hornblende and biotite ages outof shear bands date deformation and rapid coolingthrough the middle crust at 30–28 Ma.

Chiquimula batholith within Las Ovejas

complex

Coarse-grained granodiorite 5C-13 is part ofthe Chiquimula batholith with the mineral assemb-lage quartzþK-feldsparþ plagioclaseþ biotiteþhornblende. Subhedral and anhedral plagioclasecrystals are typically altered to sericite and haveAn-contents of 35–55 mol% with normal zoningpatterns. Twinning is in part deformational. Sub-hedral orthoclase has a K-feldspar componentof 80–90 mol% and displays perthitic unmixing.

Prismatic amphibole is magnesiohornblende (classi-fication of Leake et al. 1997) with relicts of augite inthe cores. Reddish-brown biotite contains up to4.6 wt% TiO2, has an XMg of c. 0.45 and is partlytransformed to chlorite. Accessories are opaqueminerals, apatite, and zircon. Anderson & Smith’s(1995) Al-in-hornblende geobarometer yields100–180 MPa at 740–860 8C, calculated with theTi-in-hornblende and plagioclase–hornblendegeothermometers (Colombi 1988; Blundy &Holland 1990; Holland & Blundy 1994). OurAr/Ar geochronology demonstrates that the Chiqui-mula granite is a composite batholith (c. 90 Mabiotite–chlorite age, 5C-11; c. 20 Ma K-feldspar,5C-13); assuming that the 20 Ma K-feldspar ageof 5C-11, adjacent to 5C-13, is close to the intrusionage of this shallow granodiorite (3–5.5 km), theaverage exhumation rate is 0.15–0.27 mm/annum.This supports our inference (see above) that rela-tively rapid exhumation in the Las Ovejas com-plex is pre-20 Ma.

Interpretation of new petrology

Garnet amphibolite and felsic gneiss of the centralChuacus complex in the El Chol area preserve, inaccordance with Ortega-Gutierrez et al. (2004), ahigh-P event. Garnet gneiss 5CO-4b probablyrecords part of a prograde path (c. 580 8C atc. 2 GPa) that culminated at c. 675 8C andc. 2.4 GPa (5CO-5a). Isothermal decompressionto c. 660 8C at c. 1.3 GPa followed the near ultra-high pressure conditions and the retrograde pathcan possibly be traced to c. 520 8C at c. 0.7 GPa.Only restricted areas of the Chuacus complexescaped the regional overprint at 450–550 8C and0.7–0.8 GPa. The age of the high-P event ispre-238–218 Ma, the age of two-mica orthogneissintrusion (U–Pb ages) and high-T metamorphism(Ar/Ar ages) in the El Chol region. The high-Pevent is probably post-638–477 Ma, the lowerintercept U–Pb zircon age (including errors) ofhigh-P garnet–K-feldspar gneiss at location5CO-4. It is possibly older than the c. 302 Malower intercept U–Pb age of a leucosome inter-preted to be associated with eclogite decompression(Ortega-Gutierrez et al. 2004); we have so far notbeen able to reproduce this age at the same localityusing SHRIMP geochronology (unpublishedresults). The high-P event should also be youngerthan c. 440 Ma, the age of the youngest zircons inChuacus paragneiss (Solari et al. 2008). Thec. 403 Ma age of likely metamorphic zircon incorpo-rated in the Triassic two-mica orthogneiss indicatesDevonian metamorphism in the Chuacus complex.The regional �450 8C and 0.7–0.8 GPa overprintis Late Cretaceous and the most outstanding featureof most of the Chuacus complex. This regional


pressure-dominated metamorphism is associatedwith local magmatism. The PT estimates of theCretaceous event are grossly similar in the Rabinalcomplex of western Guatemala and the Chuacuscomplex of central Guatemala, but somewhathigher temperatures and pressures occur in centralGuatemala. The microdiorite from the Juan de Pazultramafic sheet of the North El Tambor complexprobably records seafloor hydrothermal overprintfollowed by low-grade Cretaceous metamorphism.

Our reconnaissance petrology of the northernChortıs Block identifies Cenozoic amphibolite-facies metamorphism, migmatization, local magma-tism and associated deformation for the Las Ovejascomplex starting at �40 Ma. Peak PT conditionsare c. 650 8C at 0.7–0.8 GPa, veiling earlier high-grade metamorphism associated with magmatismand migmatization that is Permo-Triassic and Juras-sic (273–245, 166–158 Ma). Distinctly different,Middle Jurassic (c. 155 Ma) prograde, lower green-schist- to epidote–amphibolite-facies metamorph-ism is preserved in the non-migmatitic Sanaratecomplex of southern Guatemala, west and outsideof the Cenozoic-shaped Las Ovejas complex. Thelatest Cretaceous–Cenozoic Chiquimula batholithis a composite, shallow crustal level intrusion(3–5.5 km). Sub-volcanic c. 20 Ma intrusives inthis batholith demonstrate exhumation fromc. 25 km during 40–20 Ma, supporting a two-stagedeformation–exhumation history within the LasOvejas complex; that is relative rapid pre-20 Mafollowed by slower post-20 Ma exhumation (seeabove).

Figure 8a summarizes the available P–T– t datafrom the Rabinal, Chuacus, Sanarate and Las Ovejascomplexes. The intricate metamorphic evolutionin the Chuacus complex apparently is relatedto Silurian–Carboniferous near ultra-high pressuremetamorphism and its retrogression into theepidote–amphibolite-facies field; it is overprintedby Triassic magmatism and metamorphism, andregional Cretaceous epidote–amphibolite toamphibolite-facies reheating. The northern ChortısBlock contains crust dominated by Jurassicepidote–amphibolite metamorphism, only weaklyoverprinted in the Cenozoic; this Cenozoic meta-morphism defines the Las Ovejas complex, whichis characterized by epidote–amphibolite toamphibolite-facies reheating. Figure 8b comparesthe P–T– t evolutions of the melange complexes(North and South El Tambor) and the continentalblocks onto which they were emplaced. Jurassic–Early Cretaceous North El Tambor complexoceanic and immature arc rocks incorporatedc. 131 Ma eclogite into an accretionary wedge thatreached crustal levels in the Late Cretaceous(c. 76 Ma) and was incorporated as the roof nappeinto the Chuacus complex–southern Maya Block

imbricate stack that initiated contemporaneously(c. 75 Ma). Jurassic–Cretaceous South El Tamborcomplex oceanic rocks incorporated c. 132 Ma eclo-gite into an accretionary wedge, reached crustallevels at c. 120 Ma, and were emplaced ontoChortıs-Block crust, whose last thermal overprintwas during the Middle Jurassic. The South ElTambor complex was variable affected by theCenozoic deformation and metamorphism thatcharacterizes the Las Ovejas complex (see below).Characteristically, the South El Tambor complexis preserved outside the area of strong Cenozoictectono-metamorphic overprint, i.e. outside thecentral and eastern Las Ovejas complex; we mayspeculate that parts of the latter may actually beSouth El Tambor complex subjected to Cenozoichigh-grade metamorphism and deformation.

Structure and kinematics

Rabinal complex and Late Palaeozoic cover,

western Guatemala

Figures 9 and 10 compile structural and kinematicfield and laboratory data and Figure 11 shows repre-sentative deformation features and dated synkine-matic minerals. The Rabinal complex in westernGuatemala shows regional, heterogeneously distrib-uted, low-grade deformation (D2) that locally over-prints high-grade, partly migmatitic flow structures(D1). Most pronounced is a deflection of the firstfoliation, s1, into ubiquitous, sub-vertical, sinistral,D2 shear zones. In this study, we concentrated onthe Late Cretaceous (see below) greenschist-faciesevent D2; its low grade is distinct, as the sameevent reaches amphibolite facies in the centralChuacus complex of southern Guatemala and iscoeval with magmatism. Another distinct map-scalefeature is a NW structural trend (foliation, stretchinglineation, shear planes), oblique to the c. east trendof the Polochic fault zone; the latter is clearlyyounger (see above).

The basement mainly comprises locally migma-titic paragneiss and orthogneiss, amphibolite and(hornblende)–epidote schist, augen gneiss of mag-matic protolith, pegmatite and variably strainedgranitoids. The metagranitoids also intrude possiblyPrecambrian, low-grade sedimentary rocks (SanGabriel unit of Solari et al. 2009). We also studiedlocalities of Late Palaeozoic metasediments (oftenconglomerate; Chicol and Sacapulas Formations).D1 shows high-T plasticity, dominant coaxial flat-tening (s fabrics), and intrafolial and isoclinalfolds, which refold an even earlier, relict foliation.D1 is accompanied by local migmatization, tracedby feldspar layers. D1 quartz LPO shows medium-to high-T prism ,a. glide (G20s of stop G45;Fig. 9, bottom right) that contrasts with D2 low-T


Fig. 10. Additional structural data from the Maya Block of western Guatemala. Legend for structural symbols as inFig. 9. ccw, counterclockwise.


Fig. 10. (Continued) .


Fig. 11. Representative dated minerals and deformation-geometry and kinematic structures. (a) Chuacus complex and Rabinal complex rocks. 1, Stop G27, sample G7s: weathered retrogradebiotite gneiss with sinistral brittle–ductile faults and pull-apart-type tension gashes that were locally filled with pseudotachylite (c. 70 Ma Ar/Ar whole rock). 2, Western Guatemala: stopG46, sample G21s, partly mylonitic epidote schist; epidote and feldspar are brittle, quartz is ductile with subgrain-rotation recrystallization; quartz layers show synrecrystallization s–c fabricand strain concentration in shear bands (sb); phengite (c. 0.8 GPa) dates deformation at 67 + 1 Ma (stop G1, Ar/Ar); inset (top-left, stop G8, sample GM8s) shows bulging recrystallization inquartz of mid-Jurassic granitoid that is locally mylonitized along discrete Cretaceous low-temperature shear zones. 3, Stop 5CO-6: Triassic orthogneiss with pegmatite dykes and podsdeformed top-to-NE during amphibolite-facies, c. 74 Ma (U–Pb zircon) flow. 4, Stop 854: second deformation, likely to be Cretaceous, sinistral shear zones anastomose around relict garnetand white mica microlithons. 5, Stop 856: undeformed pegmatite dykelet (quartzþ plagioclase, plagioclase mostly converted to epidote) intrudes parallel to c planes of second deformation.Recrystallized s–c fabric in quartz-rich layers of orthogneiss; pegmatite is Late Cretaceous (e.g. samples 885, 5CO-6b). Feldspar in orthogneiss is broken but locally recrystallized. 6, Stop5PB-9, eastern Guatemala: high-strain, sinistral shear zone dated by synkinematic white mica at c. 76 Ma (Ar/Ar and Rb–Sr). (b) Chortıs Block rocks. 1, Stop 5PB-5, Las Ovejas complex:synkinematic amphibole (hastingsite) and biotite in c. 650 8C, 0.75 GPa (garnet) amphibolite date rapid cooling at c. 30 Ma. 2, Stop G24, sample G5s, Sanarate complex: c. 155 Masynkinematic amphibole and chlorite in amphibolite. 3, Stop 5C-26, Las Ovejas complex: weakly deformed, migmatitic amphibolite, hornblende is synkinematic at c. 35 Ma (Ar/Ar) withslow cooling to or reheating at c. 20 Ma (Ar/Ar hornblende and biotite). 4, Stop 5C-2, Las Ovejas-complex amphibolite: c. 20 Ma synkinematic hornblende in mylonite. 5, Stop 5C-37, LasOvejas complex: retrograde amphibolite–biotite gneiss sequence with mostly foliation-parallel leucogranite dykes and quartz-segregation veins. Part of the high strain zone (bottom) anddetails (insets top) shows sub-vertical, greenschist-facies, second deformation under sinistral wrenching. Hornblende dates beginning of deformation at c. 39 Ma and biotite retrogression at c.20 Ma (Ar/Ar). Hammer and coins for scale. 6, Stop 5H-4, Las Ovejas complex, Honduras: greenschist-facies, sinistral shear in diorite; hornblende and biotite date deformation between 38and 25 Ma (Ar/Ar). (c) Northern foreland fold–thrust belt. Stop G60: Cretaceous limestone; NE–SW shortening and NW–SE extension indicated by pressure-solution surfaces, fold-axisparallel tension gashes, and faults. Ductile strain is recorded by elliptical shape of crinoid stems; foliation dips steeply into the picture, bedding is sub-parallel to the outcrop surface. Hammer,compass, and coin for scale.

Fig. 12. Overview map (top left) locates geological–structural map (bottom right) in the Maya Block of central Guatemala. Main structural features are plotted both on map and in stereonets (lower hemisphere, equal area). See Figs 9 & 10 for legend; B,fold axis.

quartz LPO (see below). Quartz is coarse-grainedand recrystallized by grain-boundary migration.D1 is Late Triassic (c. 215 Ma), suggested by ourU–Pb zircon age of phyllonitic (due to D2) migma-titic gneiss (G18s of stop G1). D2 formed 10–80 cmwide, chlorite- and carbonate-bearing shear zonesthat boudinaged s1 (e.g. G1), imposed locally well-developed augen- to ultra-mylonites that evolvedduring decreasing temperature into ductile–brittlemylonite with clinozoisite-, epidote- and quartz-bearing faults (e.g. G3) and late fracturing (e.g.G45). These discrete shear zones widen into beltsof mylonite, transforming the basement gneissesinto biotite–chlorite schist and epidote–hornblendephyllite. Strain is locally prolate (dominantl-fabrics; e.g. G56/57). D2 quartz LPO indicatesbasal ,a. slip (G1) with a strong coaxial flowcomponent (c-axis cross-girdles). Quartz recrystal-lization mechanisms are variable but typicallylow-T and dynamic; subgrain rotation is dominant(e.g. G21s of stop G46; Fig. 11a-2) but evolvesout of bulging recrystallization in quartz ribbons(e.g. G25s of stop G56, Figs. 11a-2, inset). Horn-blende and feldspar are broken. Plagioclase showsrare incipient stages of recrystallization along twinplanes; thus, deformation occurred mostly belowc. 450 8C, evolving to very low grade (�300 8C).Sinistral shear criteria are ubiquitous: flow folds,s clasts, shear bands and s-c- fabrics in layersof quartz recrystallization (Fig. 9). Epidote(+hornblende) schists G19s (stop G1) and G21s(stop G3) provide quantitative PT estimates for theinitial, higher-T stages of D2 at 450–500 8C and0.7–0.8 GPa. D2 is precisely dated at 67 + 1 Maby synkinematic white mica (G19s of stop G1,Fig. 11a-2). D2 also developed in the Palaeozoicmetasediments and is more pronounced (sinistral)transpressional (e.g. G4, Fig. 9) than in the under-lying basement. Location GM36 shows stretchedconglomerate (Fig. 9) cut by aplitic veins with thesame c. east–west extension that stretched thepebbles; low-T plasticity with dynamic subgrainrotation occurred in quartz clasts. The basementrocks are variably overprinted by faults that arerelated to the Polochic fault zone (Fig. 10); the prin-cipal faults strike WNW (Polochic-parallel) andnorth. Faulting is �20 Ma, given by the c. 20 MaU–Pb zircon lower intercept of samples GM13sand GM14s (stops GM42, GM43) and the c.23 Ma AFT age of stop GM35 (sample GM9s); itis best constrained by the TFT and AFT agesalong the Polochic fault (11.5–4.8 Ma, see above).

Rabinal complex, Rabinal granitoids and

low-grade metasediments, central Guatemala

The Rabinal complex in central Guatemalacomprises paragneiss that typically resembles

metagreywacke (coarse-grained feldspar and whitemica) and K-feldspar porphyroblastic granitoid(U–Pb zircon ages at 485–410 Ma, stop 5CO-1)that were transformed to quartz phyllite and augengneiss. The distinction between ortho- and para-gneiss is problematic at many locations and makesmapped lithological boundaries questionable (e.g.BGR 1971, Fig. 12). Locally marble and amphibo-lite intercalations occur. The major deformation,attributed to the WNW-trending Baja Verapazshear zone by Ortega-Gutierrez et al. (2004), islocally mylonitic and low grade. Foliation dipsintermediate to steeply south(west) and the stretch-ing lineation (str1) plunges intermediate to the SW(Figs 12 & 13). A continuous kinematic evolutionfrom ductile shear bands to ductile–brittle faultsand quartz-filled, en-echelon tension gashes (e.g.stops 5CO-1, 5CO-3, 867) indicates sinistral trans-pression. At station 5CO-1, typical Rabinal augengneiss, shear bands evolve into ductile–brittlefaults (quartz ductile), and faults with largequartz–white mica fibres that are conjugate tokink bands that are partly faulted and indicatenorth–south shortening and east–west extensionalong str1. Quartz shows typical low-T plasticitywith subgrain-rotation recrystallization. K-feldsparis brittle. Local ultramylonite (,1 m thick, stop867) has cogenetic chloriteþ sericite and quartzribbons that show weak bulging- and dominantsubgrain-rotation recrystallization. Late kink foldsare interpreted as flow heterogeneities at the endof deformation. Quartz fibres in the strain shadowof pyrite are un-recrystallized. The post-245 Ma(see above) phyllite and conglomerate (mostlyintermediate volcanic rock pebbles, stops865–866) records identical, low-T structural geo-metries and kinematics. Paragneiss (stop 867) pre-serves pre-Cretaceous deformation with largewhite mica and pre-existing foliation and felsicveins that are cut by orthogneiss that is typicalRabinal gneiss.

Chuacus complex, central Guatemala

Central Guatemala exposes the deepest part of theChuacus complex studied by us (Figs 12 & 13) andallows insight into pre-Cretaceous deformation–metamorphism–magmatism. The Cretaceous struc-tural overprint increases both toward south andnorth, i.e. the North El Tambor complex of thehanging wall. Typical Chuacus-complex rocks aregarnet-bearing para- and orthogneiss, garnet amphi-bolite that is partly retrograde eclogite (see petrol-ogy), and marble (e.g. stops 880–884, 886,5CO-4–6). Felsic and mafic rocks are often concor-dantly interlayered and resemble a bimodal volcanicsequence dominated by metasediments. The layer-ing is accentuated by felsic sills (dominant) and


dykes that crop out as partly pegmatitic orthogneissand locally comprise 50% of the rock volume. Ournew geochronology assigns them to the Triassic(see above). This rock association is sheared,folded, and thermally overprinted by syn- to post-tectonic feldsparþ hornblendeþwhite micaþ

garnet (+kyanite in quartz veins) growth. Local,large-scale fluid flow is indicated by up to 1 mthick quartz-segregation veins and ubiquitousalbite growth. Locally, but increasingly abundantto the south, a second type of pegmatite is observed(e.g. stop 884). It is muscovite-rich, crosscuts

Fig. 13. Additional structural data from the Maya Block and the Early Cenozoic Subinal Formation in centralGuatemala. Legend for structural symbols see Figs 9, 10 & 12.


the sills/dykes, and is associated with loweramphibolite- to upper greenschist-facies meta-morphism. These pegmatites are undeformedwhere massive but sheared along quartz-richlayers. They are Late Cretaceous as indicated byc. 74 Ma metamorphic overprint in the host rocks

(U–Pb zircon, stop 5CO-6) and their 74–62 Mawhite mica ages (see above).

Where preserved, Triassic deformation is dis-tinct. Foliation is sub-vertical and folded tight-isoclinally to ptygmatically with axes parallel tothe mineral stretching lineation (F1 ,a. type





folds) that trends NNW (Fig. 12, stops 5CO-5,880–886) and thus is clearly different from theoverall trend of the Late Cretaceous deformation(c. ENE). F1 verge mostly (north)east and arelocally sheath folds. At these locations, overprintingCretaceous F2 folds are steeply plunging, open,northeast-vergent, and associated with dextraland sinistral (conjugate) shear bands (Fig. 12).Garnet–K-feldspar gneiss (5CO-4), resemblingRabinal augen gneiss, has old garnet with an internalfoliation that is 908 to the external (Cretaceous) foli-ation. Overall, Triassic deformation appears coaxialand the structures are annealed (widespread albiteovergrowth, see above).

Structural geometry and kinematics of theregionally dominant Cretaceous deformationappear simple (Figs 11a-3, 12 & 13): the stretchinglineation consistently trends c. ENE and parallelsthe ubiquitous B2 axes; where s1 is sub-vertical, itis transposed by sinistral shear zones/bands;where s1 is sub-horizontal, shear zones/bandsrecord sinistral transpressive, top-to-NE flow. Latefaults have huge fibres and indicate north–southshortening (e.g. stop G29). This Cretaceous high-strain, NE–SW stretching is suggested by variousfeatures. Stop G31, for example, shows relativelyidiomorphic, inclusion-free, pre-Cretaceous garnetthat is fractured and connected with chlorite; asecond generation of garnet and chlorite is co-genetic. Quartz shows transition from subgrain-rotation to grain-boundary-migration recrystalliza-tion (e.g. stop G31). At stops 880–886 late,mostly unfoliated/unfolded dykes/veins areexposed; some of these quartzþ feldsparþmuscoviteþ biotite pegmatites show high-strainboudinage. At stop 887, a several metres thick,weakly foliated pegmatite crosscuts the major(Triassic) foliation but deflects it sinistrally.F3 refold F2 openly. Where F3 is not developedl-tectonites are common. Our interpretation is thatfolding is represented by prolate strain (e.g. stop887). Our geochronology constrains coolingfrom c. 550 to 275 8C and thus deformation at75–65 Ma (see above).

Stations 854 and 856 reveal Chuacus complexthat is dominated by orthogneiss (Fig. 12). At 856,,0.5 m thick, D1 mylonite–ultramylonite shearzones cut augen gneiss, biotite granite and aplitedykes. The host rock is weakly deformed at high-Tpreserving NW trending str1. Discrete ultramyloniteshear zones that differ in degree of localization – theyounger are more strongly localized – overprint afirst, east striking, dextral, high-T shear zonefabric. D1 is syn-migmatitic, F1 are isoclinal. D2

imprints ENE-trending ultra-mylonitic–mylonitic,Ramsay–Graham-type shear zones with late chlor-ite–quartz veins, in which quartz is fibrous tracingthe NW–SE str2. Quartz and plagioclase in the

sinistral, often anastomosing shear zones(Fig. 11a-4) show subgrain-rotation recrystalliza-tion and reaction softening by transfer to sericiteand epidote, respectively; locally feldspar is recrys-tallized (c. 500 8C). At 854 and 856, undeformedpegmatite and aplite dykes/dykelets cut migmatitic,sillimanite- and white mica-bearing ortho- and para-gneiss and are younger than the D2 fabric(Fig. 11a-5). Pure quartz layers show prism ,a.slip and subgrain-rotation with transitions to grain-boundary migration recrystallization. Common F2

refold both s1 and D2 shear zones, imposingprolate strain. Radiometric age control on thesedeformations is weak (854F); migmatization andD1 are pre-149 Ma, however, structural geometry,kinematics and associated metamorphism suggesta Late Cretaceous age for D2.

We also studied the Chuacus complex along asection at its easternmost outcrop at the base of theJuan de Paz ophiolite unit (Fig. 12). The structuraltrend changes to ENE, from WNW in western andeast in central Guatemala. Deformation is continu-ous from greenschist facies to brittle (boudinagedquartz rods) and heterogeneous. At station 5PB-9,deformation is high-strain, mostly coaxial with sinis-tral and dextral shear bands and ,5 cm thick mylo-nite zones, shows coaxial, basal ,a. slip in quartz,and open ,a. folds. The whole rock–synkinematic white mica (Fig. 11a-6), Rb–Sr iso-chron at 76.5 + 0.6 Ma precisely dates deformation.At station 5PB-10 (Fig. 12), ubiquitous vertical,sinistral shear bands, a single- to type-I cross-girdlequartz c-axis orientation and strong preferred orien-tation of a-axis and m-planes indicate dominantrhomb and prism ,a. slip under somewhat higherdeformation temperatures than at station 5PB-9and dominant non-coaxial flow.

Metasomatized southern Chuacus complex

rocks below the North El Tambor complex,

central Guatemala

The immediate footwall of the North El Tamborserpentinites comprises imbricates that stackChuacus-complex rocks, Chuacus rocks with low-grade sedimentary rocks (Jones Formation and SanLorenzo marble), and these rocks with serpentinite(Figs 12 & 13). Rocks that are imbricated are phyl-lite (retrograde biotite gneiss), epidotized biotitegneiss, (+garnet)–actinolite–clinozoisite–chloritephyllite, marble and low-grade siltstone. Theserocks show neocrystallization of calcite, quartz,Mg-chlorite and albite (e.g. stops 5PB-15, 860)depending on the surrounding lithology. The bound-ary between serpentinite and Chuacus-complexrocks is, where exposed, a shallow-dipping brecciazone with serpentinite and marble clasts (stop


G26, Fig. 13). Nearly all of the stations (Fig. 13)show relict, mostly sub-horizontal, high-grade foli-ation (s1) that is folded (e.g. G26, G27, 5C-36).During D2 sinistral ductile–brittle shear bandsdeveloped where s1 was steeply dipping (e.g.5PB-15, 860). Stretching is WSW–ENE and exem-plified by s1 boudinage. Quartz-filled tension gashesare ubiquitous (e.g. G26, 5PB-15) and testify to fluidflow. F2 flexural-glide folds are north(east) vergentand show lengthening along B2 (e.g. G26, 5C-36).Quartz deformed by subgrain-rotation recrystalliza-tion (e.g. G26) and basal ,a. glide (5C-36). TheLPO indicates dominantly coaxial flow. Whitemica at stop G26 is sericite and its Ar/Ar agedates deformation at �60 Ma. This corresponds tothe Rb–Sr white mica–whole rock isochron at c.63 Ma from a chlorite phyllite. Faulting started in

the ductile–brittle regime with a dominant northeaststriking set. At station G27 pseudotachylite (Figs11a-1 & 13, sample G7s) fills pull-apart tensionfractures. These are again involved in continuousdeformation, as east-trending quartz fibres formedas strain shadows. The pseudotachylite dates D2 atc. 67 Ma.

North El Tambor complex on Chuacus

complex, central Guatemala

We analysed imbrications of serpentinite, ophical-cite, and felsic volcanic rocks (stop G34) andserpentinites (stops G28, G35, G38, GM11-13,5PB-7; Figs 13 & 14). All stations show a low-grade, ductile–brittle fabric with shear bandsand faults that slipped sinistrally, coeval with

Fig. 14. Geological map (bottom) of the Motagua suture zone with structural data from the northern Chortıs Block, theNorth and South El Tambor complexes, and the Sanarate complex; the outline of the latter is conjectural. In additionstructural data of Cretaceous to Late Cenozoic sedimentary and magmatic rocks straddling the suture are shown. Mainstructural features are plotted both on map (this figure) and in stereonets (lower hemisphere, equal area, Fig. 15). SeeFigs 9 & 10 for legend. Adapted from Bonis et al. (1970).


shortening indicated by folding and oblique-slipfaulting. White mica ages (c. 65 Ma) of greenschist-facies rocks (Chuacus complex, stops G26, G29, seeabove) in the immediate footwall of these ElTambor rocks suggest latest Cretaceous defor-mation. At several stations, tectonic melangeoccurs along major north dipping detachments thatcontain foliated, moderately serpentinized ultra-mafic rocks (G34, G38) and pure serpentinite(G35; Figs 13 & 14) with brittle–ductile s–cfabrics. The ultramafic bodies are surrounded byan anastomosing, lenticular, scaly foliation thatrecords overall coaxial deformation. Shortening isNNE–SSW and extension is WNW–ESE. Moststations contain a younger fabric with unclearage relations: sinistral shear and normal faultingboth with NW–SE extension. Late-stage, flattening-type extension is possibly caused by topographiccollapse.

Northern foreland fold–thrust belt,

Guatemala

The ‘Laramide’ foreland fold–thrust belt north ofthe Polochic fault zone and the overlying North ElTambor complex allow comparisons of structuralgeometries and kinematics with those recorded inthe Chuacus basement in the south. The strati-graphic range observed at our structural stationsranges from Permian to Upper Cretaceous, that isPermian Chochal and Tactic Formation limestoneand shale, and Lower and Upper CretaceousCoban and Campur Formation dolomite and lime-stone (BGR 1971). In the following, we give anoverview but emphasize that the stereoplots(Figs 10 & 13) contain a cornucopia of structuraldetail, for example, geometries, deformationevolution and changes in the palaeostress field.Deformation age is difficult to constrain directly.Structural and kinematic compatibility with thewell-dated deformation in the Chuacus complex(see above) and that in the Sierra Madre Oriental–Yucatan fold belt (e.g. Gray et al. 2001) suggestthe major deformation, sinistral transpression, islatest Cretaceous–Early Cenozoic.

In western Guatemala, stations G59–G61 withinCretaceous limestone and minor shale recordfolding with well developed ductile–brittle, NW–SE extension along the regional fold trend(Figs 10 & 11c-1). Deformation starts during layer-parallel shortening before buckling with verticalextension by tension gashes (tgold) and continueswith buckling and layer-internal ductile flow withfoliation development, foliation boudinage and cre-ation of a fold axis-parallel stretching lineation.Later stages develop conjugate but mostly sinistralfaults and oblique stylolites that show strongvolume-loss by pressure solution; dissolved

material is partly deposited in veins that compriseup to 20 vol% of the rock. Fold axis-parallel exten-sion is accentuated by a second set of tension gashes(tgyoung) and conjugate faults that represent tiltedhorst-graben structures; the latter exactly constrainextension at 1288 (inset stop G61–G62, Fig. 10).Abundant crinoids in the limestone record up to50% shortening due to volume reduction that ismostly accommodated by bedding-parallel stylo-lites (Fig. 11c-1). Other stations, both in limestoneand serpentinite, show less ductile flow but morepronounced faulting. Folding is mostly gliding inbedding; the slip direction is oblique to the foldaxes (e.g. stop G55). Common to most stations isrotation of s1 from east–west to north–south intime and sinistral transpression.

In central Guatemala, Permian limestone and thecontact zone between Cretaceous limestone and ser-pentinite provide a similar albeit less completedeformation history. Sinistral transpression androtation of s1 imply increased shortening and lessstrike–slip over time (e.g. G43, Fig. 13). Theweak structural record of serpentinite emplacementand the overall paucity of shallowly dipping faultsare notable. Early fault sets indicate c. northeast–southwest shortening along conjugate strike–slipfaults (e.g. G39). At station G40, c. north–southshortening along mostly strike–slip faults formedanastomosing fault zones with boudinage andeast–west extension. Normal faults may be topogra-phically induced, and late cataclastites likely recordactive sinistral wrenching along a strand of thePolochic fault zone. The normal faulting observedat station G41 could also be gravitationallyinduced or, alternatively, a transtensional continu-ation of sinistral wrenching.

Cenozoic volcanic rocks and Early Cenozoic

Subinal-Formation red beds, western and

central Guatemala

In western Guatemala, Cenozoic pyroclastic rocks,welded tuffs and, locally, sedimentary rocks coverthe Rabinal complex and are exposed along thePolochic fault zone. Our data (Figs 9 & 10) highlightthree strike–slip dominated events. From older toyounger these are: (i) c. north–south shortening;(ii) c. NE–SW shortening, mostly along c. easttrending sinistral strike–slip faults. This is by farthe dominant set and traces the Polochic fault-zonedeformation. The strata dip up to 808, indicatingstrong local shortening. A similar stress field butdominated by normal faults also occurs in subhori-zontal pyroclastic rocks that likely are part of theCenozoic ignimbrite province of Central America(c. 15 Ma; e.g. Jordan et al. 2007); we thus attributethis major event to the currently active stress field.(iii) Normal faulting with widely dispersed slip


directions. This event probably records topographiccollapse.

Deformation of the Subinal-Formation red bedsstarted with c. north–south shortening that tiltedbeds (open to tight folds) and produced a fracturecleavage (Figs 13 & 14). At stops 872–873, twosets of older faults have constant ENE–WSWextension (s3) but s1 and s2 permutated. Thestrike–slip faults have oblique lineations so thatfaulting probably started during late stages offolding. Faulting recorded at all stations has aclear sinistral-transpressive component. Latenormal faulting is interpreted as prolongation offolding-related deformation with extension alongthe fold axes becoming dominant.

Sanarate complex, central Guatemala

The Sanarate complex, the westernmost basementoutcrop south of the MSZ, comprises retrogradeamphibolite, chlorite–epidote–hornblende schist,micaschist and chlorite phyllite with peak PT con-ditions at c. 550 8C and c. 0.65 GPa. The basementis in fault contact with Cenozoic andesite and redbeds (stops G23, G24, 5G-1; Figs 14 & 15). Foli-ation and relict isoclinal folds constitute D1. D2

shows low-T, partly mylonitic quartz–white micafabrics along an east trending, vertical foliation;shear sense is sinistral. Quartz is post-tectonicallyannealed. D3 occurs both in the red beds/volcanicrocks and the basement. ENE–WSW shorteningthrust the basement in a sinistral-oblique senseonto the red beds. Two generations of hornblendeexist in the studied amphibolite. The dominant syn-kinematic younger generation (Fig. 11b-2) and alsosynkinematic white mica in the micaschist/phyllitedate D2 as Middle Jurassic (samples G3s, G5s,stations G23, G24, Fig. 5e). The reheating (quartzweakly annealed) is probably Cenozoic (low-Tsteps in G5s). Adjacent stations 853 and 870–871comprise massive amphibolite and rare garnetamphibolite and leucogabbro close to the boundarywith the South El Tambor melange rocks. Theirapparently lower-grade metamorphism disting-uishes them from the Las Ovejas complex amphibo-lites. Asymmetric boudinage of hornblende–quartzveins, s clasts and late biotite-rich shear bandsassociated with kink bands record top-to-NE flowwith a slight dextral component along westdipping foliations. Tight to isoclinal flow folds arepresent in high-strain zones (Fig. 15). This defor-mation remains to be dated.

Las Ovejas complexes and San Diego phyllite,

Guatemala and western Honduras

Stations 5C-2, 5C-28, 5C-26, 27 and 5C-22–24 arein basement lacking retrograde metamorphism of

the Las Ovejas complex south of the centralMotagua valley and north(east) of serpentinite andmelange rocks of the South El Tambor complex(Figs 14 & 15). This area comprises biotite granitewith diorite xenoliths, gabbro-amphibolite, migma-titic biotite gneiss and garnet-bearing leucosome.The main deformation is of variable intensity,from weakly foliated to locally mylonitic–ultra-mylonitic but constant in orientation. Foliation hasintermediate dips toward the NW, kinematicallyrelated shear bands trend NE and dip, in general,steeper than the foliation, and the stretching linea-tion is sub-horizontal NE–SW, parallel to axes of,a. type folds. These structures overprint relictearlier deformation fabrics and began syn-migmatic(Fig. 11b-3) with melt locally concentrated alongshear bands; shear is overwhelmingly sinistral.Dykes intrude parallel to the foliation (5C-22, sub-volcanic granite), mostly normal to the stretchinglineation (5C-2, diorite), thus indicating syntectonicmagmatism or crosscut irregularly (5C-28, pegma-tite). Synkinematic hornblende and white micadate deformation as pre-to syn-35 Ma; a cluster ofc. 10 Ma younger ages (mostly cooling through300 8C) indicates later structural reactivation, assuggested by locally well-defined, synkinematicmineral fabrics (e.g. stop 5C-2, Fig. 11b-4).

Paragneiss and micaschist at stop 5C-3 boundSouth El Tambor serpentinite and show syntectonic,greenschist-facies retrograde metamorphism. Latesinistral faults juxtapose these rocks with the ser-pentinite. Stations 5C-37 and 38 at the southeasternedge of the same South El Tambor allochton com-prise migmatitic biotite gneiss, amphibolite andleucogranite and their retrograde products. D1

synamphibolite-facies deformation with local mig-matites is overprinted by localized but pervasivegreenschist-facies D2 (Fig. 11b-5). The biotitegneiss and the amphibolite are in part extremelydeformed and often l-tectonites; s1 and sinistral-transtensional shear bands dip intermediate to NWand str1 plunges shallowly. The tourmaline–garnet–white mica leucogranites postdate the mig-matitic gneiss, intrude mostly as sills, and arelocally deformed synintrusive with high strain. D2

induced vertical sinistral shear zones, ,a. folds,foliation boudinage and syn- and mostly antitheticfaulting of leucogranite dykes and quartzsegregation veins (Fig. 11b-5). The transition fromshallowly dipping D1 shear bands to vertical,ductile–brittle D2 ones is abrupt (within c. 10 m).Hornblende (amphibolite) and white mica (leuco-granite) date the initiation of deformation andsyntectonic intrusion at 40–35 Ma. The likely kin-ematic evolution from ductile (�40 Ma) toductile–brittle (c. 20 Ma, biotite) could be inter-preted in this outcrop as long-lasting progressivedeformation instead of structural reactivation (see


above). Stations 5PB-1–5 and 861, 862 exposeunretrogressed Las Ovejas complex migmatiticbiotite gneiss, intercalations of (garnet) amphibolite,biotite–white mica–garnet–tourmaline pegmatite,marble and calcsilicate rocks. Locally these rocks

are post-tectonically annealed due to their proximityto granitoids (Fig. 14). At location 5PB-5 massivemigmatite is enclosed in granite. Garnet amphibolitecontained in the migmatite yielded c. 650 8C at 0.7–0.8 GPa peak metamorphic conditions and preserves

Fig. 15. Structural data from the Chortıs Block and Cretaceous to Late Cenozoic sedimentary and magmatic rocksstraddling the Motagua suture zone. Legend for structural symbols as in Figs 9, 10 & 12.


pre-migmatitic garnet relicts. Deformation startedprior to migmatization, is locally high-strainwith migmatitic biotite–gneiss mylonite, and issinistral-transpressive with local s–c mylonites. Itcontinued with sinistral ductile–brittle faulting.

Migmatization was at c. 36 Ma (U–Pb zircon andAr/Ar hornblende, see above). Station 5C-15shows the Jocotan fault as a ductile–brittle sinistralfault zone. North–south ductile shortening withwell-developed sinistral shear bands changed to



NE–SW compression in the brittle field. Low-Tplasticity is recorded by bulging to dominantsubgrain-rotation recrystallization in quartz.

Las Ovejas complex of northern Honduras

Stations 5H-1–6 characterize deformation of thewestern footwall of the Sula graben in northernHonduras (Figs 14 & 15). The Las Ovejascomplex rocks and their deformation are similar tothose in Guatemala. Partly migmatitic biotitegneiss, biotite granite, diorite, micaschist andsmall white mica- and garnet- or hornblende-bearing pegmatite pods and sills deformed under

sinistral transpression that started during the mig-matitic stage. Deformation continued into theductile–brittle field with sinistral-transtensive,SE-dipping shear zones/normal faults that werelate to opening of the Sula graben. The pegmatiteseither show strain concentration with well-developed mylonite and ultramylonite or are vari-ably but mostly weakly deformed, forming irregularpods. In comparison with Guatemala, foliation isrotated, dipping moderately to steeply north withthe stretching lineation NE–SW; shear bands dipNW. Our 40–36 Ma U–Pb zircon ages from mig-matitic biotite gneiss and the c. 38 Ma Ar/Ar horn-blende age from hornblende pegmatite constrain



small-scale magmatism and deformation (seeabove). Synkinematic hornblende and biotite(Fig. 11b-6) and rapid cooling of K-feldspar ingreenschist-facies, ductile–brittle shear zonesrelated to the Sula graben suggest its initiation atc. 28 Ma. At stop 5H-7, a top-to-east shear/normalfault zone separates orthogneiss from phyllite andis likely related to the Sula graben. The orthogneissshows a relict older fabric (D1) with NE dippingfoliation. Greenschist facies with up to myloniticdeformation is present both in phyllite and ortho-gneiss. Shear bands show sinistral transpression(NE–SW shortening) with transition to brittle–ductile fabrics. Quartz shows subgrain-rotationrecrystallization. Stations 5H-8 and 9 compriseorthogneiss and migmatite with mostly Permo-Triassic U–Pb zircon ages and Jurassic–Cretaceouscooling ages. These rocks are distinctly differentfrom the Las Ovejas complex rocks along theMSZ. Migmatization is pervasive and thus differentfrom the c. 40 Ma event. D1 is coeval with migma-tization. Again, NE-dipping s1 is overprinted bydiscrete ductile–brittle shear bands/zones thatprobably are Cenozoic. Stations 5H-11 and 12comprise low-T, mylonitic orthogneiss (probably arhyolitic protolith) interlayered with thin paragneissand marble (?metavolcanic sequence). A relicthigh-T deformation, again with the characteristicNE-dipping foliation, is cut by greenschist-faciesmylonite (locally chlorite out of biotite and sericiteon mylonitic shear bands) with ,3 cm chlorite-bearing ultramylonite layers. Overall deformationis top-to-SW sinistral transpression and probablyrelated to the Jocotan fault zone.

Late Cretaceous and Cenozoic–Quaternary

sedimentary and magmatic rocks south of the

Motagua suture zone, central Guatemala

Conjugate strike–slip fault sets that indicate c.east–west shortening dominate Upper Cretaceouslimestones (stops GM4, 5 and 7; Figs 14 & 15).The quality of the field exposures does not allowus to discriminate among subgroups representing adeformation sequence or to assess block rotations.The lack of a dominant east trending set of sinistralstrike–slip faults, which we attribute to the activestress field along the Motagua fault, distinguishesthese stations from the Cenozoic–Quaternaryrocks (stops GM1, 2 and 10; Figs 14 & 15) butties them to the deformation recorded in Subinal-Formation red beds. The latest increment of defor-mation in all these stations is east–west extensionby normal faulting. Biotite granite (�37 Ma) atstations 5H-5 and 6 within the footwall block ofthe Sula graben shows brittle–ductile normalfaults and fracture cleavage related to graben

formation dated by rapid cooling at c. 28 Ma (Ar/Ar K-feldspar). Conjugate strike–slip faultingwith north–south compression was followed bynormal faulting (s3 of both events trends east) inred beds and volcanic rocks at station 5H-14within the Jocotan fault zone.

South El Tambor complex on Las Ovejas

complex and San Diego phyllite, central

Guatemala

Stations 852, 855, 864 and 869 comprise rare un-foliated blueschist, serpentinite, graphite-bearingmetasiltstone to phyllite with quartzite layers, andgreenschist of the southern South El Tambormelange. The metasiltstones show ductile–brittle,heterogeneous fracture cleavage formed bypressure solution (diffusion creep), and quartzveins in metasandstone layers. D1 shows open–tight, c. south vergent folds and thrust imbricateswith associated fault-bend folds. We do not haveage information on this melange-forming event.These rocks are cut by sub-volcanic andesitedykes and sub-volcanic granite stocks, which weinterpret as apophyses of the Chiquimula batho-lith. D2 affects the melange rocks and thesedykes/granitoids by c. east-trending (transten-sional) strike–slip and normal faults. Their low-Tnature and the presence of c. 20 Ma sub-volcanicgranite in the Chiquimula batholith (see above)suggest Cenozoic faulting.

Interpretation of new structural geology

Our work along the MSZ focused on theCretaceous–Cenozoic deformation. Clearly, how-ever, older events occurred. In the Rabinal andChuacus complexes, Triassic deformation, studiedsuperficially, shows high-T ductile flow and is dis-tinct in structural geometry and kinematics: mostly(north)west striking foliation, NW-trending stretch-ing directions, tight–isoclinal ,a. folds. Shorten-ing appears coaxial north(east)–south(west) withstrong along-strike lengthening. Associated high-Tmetamorphism, migmatization, and magmatismare Late Triassic (238–215 Ma, major group atc. 220 Ma). Foliation trajectories of the Late Creta-ceous, Maya-Block (for this discussion it includesthe Rabinal and Chuacus complexes and theircover) deformation change from WNW inwestern, to west in central, and SW in easternGuatemala. Using the (west-)northwest strike ofthe Late Cretaceous–Paleocene Mexican thrustbelt as a reference frame (e.g. Nieto-Samaniegoet al. 2006), this implies large-scale sinistral shear.The southern Maya Block in Guatemala apparentlycomprises a ductile–brittle nappe stack, whose


detailed geometry remains to be established. TheNorth El Tambor allochton in the hanging wall,itself an accretionary stack of rocks of variabletectono-stratigraphic origin and subduction depth,is underlain by an up to several hundred metres-thick carapace of imbricated rocks, containingChuacus basement and cover, altered by fluid infil-tration and metasomatism. The central Chuacuscomplex may either constitute a Late Cretaceousantiformal stack or, our preferred model, an anti-form above a blind thrust, exposing pre-Cretaceousmetamorphic and magmatic products in its core.The Baja Verapaz shear zone is the sole thrust tothis stack. The northern El Tambor sheets (BajaVerapaz, Sierra des Santa Cruz) are cut by out-of-sequence thrusts that are part of the foreland fold–thrust belt. The relatively high-P (0.7–0.8 GPa)and relatively low-T (�450 8C) metamorphism inthe Chuacus complex suggests subduction ofoceanic crust transitioning to continental crust (theChuacus geothermal gradient is �18 8C/km fromour data) and steep continental subduction; thisis supported by the nearly coeval blueschist(c. 76 Ma, North El Tambor complex) and epidote–amphibolite- to amphibolite-facies (�75 Ma,Chuacus complex) metamorphism in the accretionstack and the sinistral–transpressive deformation.Penetration of deformation and overburden (PTconditions) appear to increase from western toeastern Guatemala. Although in detail variable,meso-scale Late Cretaceous deformation is simple:north–south shortening by reverse shear zones andup to two generations of folds are accompanied bysinistral strike–slip shearing and major east–westlengthening. Tangential stretching is dominant inmost cases, as deduced from the mostly shallowlyplunging stretching lineation. In detail, overall sinis-tral transpression is typically partitioned betweenoblique stretching within the foliation and strike–slip along steeper dipping shear zones/bands tran-secting the foliation. Flow has a dominant non-coaxial component, most clearly expressed byquartz LPOs, developed during dynamic, subgrain-rotation dominated recrystallization. There is aclear kinematic continuity from ductile to brittledeformation. Although dominated by diffusioncreep, deformation is first-order identical withinthe northern foreland fold–thrust belt, the solethrust carapace below the North El Tamborcomplex sheet, and within it. Emplacement of theEl Tambor allochton, poorly constrained by ourdata, appears to change from tangential to frontalover time (strike–slip before thrusting).

Structural trends in the northern foreland fold–thrust belt and the Rabinal complex of westernGuatemala are NW and oblique to the Polochicfault zone, testifying to the younger age of thisfault. Polochic mylonitic deformation appears to

be younger than c. 15 Ma (apparent post-tectonicpluton), although our only Cenozoic age(c. 37 Ma) from western Guatemala (G27s, stopG57), from a discrete low-T mylonite zone ingneiss close to the Polochic fault zone, may hint ata Late Eocene slip history. Ductile to brittle slipalong the Polochic is dated as Late Neogene–Recent in western Guatemala (TFT and AFT ages)and at c. 8 Ma (Ar/Ar chronology) along the, prob-ably kinematically related, Tonala shear zone insoutheastern Mexico.

Cenozoic volcanic rocks record a phase of pro-nounced north–south shortening along the westernPolochic fault zone. Its active deformation is by akinematically related combination of strike–slipand normal faulting. We interpret the Subinal-Formation red beds as intra-continental molasserelated to late-stage, erosional unroofing of mostlythe Chuacus complex. Its deformation was devel-oped during north–south shortening by foldingand fold-axis parallel extension, accompanied bysinistral shear. The latest stage indicates north trend-ing graben formation.

Our study indicates that the Las Ovejas complexof the northern Chortıs Block is less a lithologicallydefined unit than distinguished from the other base-ment units by its Cenozoic metamorphism, magma-tism and deformation. It forms a westwardnarrowing belt of overall sinistral transtensionalong c. NW dipping foliations. Deformation pene-tration apparently increases toward north, the MSZ,and seems to be more localized and to die out south-ward and westward. The belt of Cenozoic shear thusappears to trend oblique to the Motagua fault.Deformation interacts with amphibolite-faciesmetamorphism, widespread migmatization andlocal magmatism, demonstrated most clearly bysmall syn- to post-tectonic pegmatites. The closespatial and temporal interaction between high-straindeformation and Cenozoic metamorphism, migma-tization and magmatism is confined to the northernChortıs Block. It appears that arc magmatism/metamorphism provided the heat input to localizedeformation. Deformation started pre-migmatiticand occurred in a kinematic continuity fromductile to brittle. Batholiths are post-tectonic. Ceno-zoic sinistral-transtensional wrenching appears,like the Cretaceous deformation of the Chuacuscomplex, partitioned between oblique stretchingalong a shallower dipping foliation and strike–slipalong steeper dipping shear zones/bands. Defor-mation began prior to 40 Ma and we suggestthat there were two major phases, �40–35 and25–18 Ma. This reactivation interpretation stillmust be tested. The Sula graben started in temporaland kinematic continuity with wrenching alongthe Las Ovejas complex at c. 28 Ma. The topo-graphically well developed active grabens,


however, are a neotectonic feature. Also theJocotan–Chamelecon fault zone was part of thisdeformation system, albeit mostly ductile–brittle.The Cenozoic event veils pre-Cenozoic tectono-metamorphic–magmatic events, with distinctstructural orientation and kinematics. The latestpreserved is Jurassic and has a NW structuraltrend. The South El Tambor complex was emplacedby south vergent thrusting and associated folding.It is involved into the Cenozoic sinistral wrenchdeformation.

The Sanarate complex with its distinct lithology,metamorphism and age of metamorphism isaffected by the Cenozoic deformation but has acharacteristic west dipping foliation and top-to-eastflow in its dominant amphibolite lithology. Thisdeformation is Jurassic or younger. Late Cenozoicdeformation, preserved in Late Cretaceous andCenozoic–Quaternary rocks on the Chortıs Block,is by distributed sinistral strike–slip. The latestevent we recognized in the Chortıs Block and theSubinal Formation is east–west extension bynormal faulting. The entire MSZ area shows gravi-tationally induced faulting due to the large topo-graphic relief.

Discussion

Here we return to the principal questions this paperaims to address: (i) which crustal sections of theMSZ accommodated the large-scale sinistraloffset along the northern Caribbean Plate boundary,when did it occur, and in which kinematic frame-work, and (ii) what Palaeozoic–Cenozoic geologi-cal correlations exist between the Maya Block, theChortıs Block, the tectono-stratigraphic complexes(‘terranes’) of southern Mexico, and the arc/subduction complexes of the northern Caribbean.First we focus on geological correlations proceedingfrom Palaeozoic to Recent, and then we refineCaribbean tectonic models for the Cretaceous andCenozoic.

Taconic–Acadian (Early–Middle

Palaeozoic) orogeny

The Chuacus complex contains a paragneiss-dominated sequence, in which metabasalts prevailamong the magmatic rocks, and quartzite andcalcsilicate/marble are minor constituents (Fig. 2d).In contrast, the hanging Rabinal complex isorthogneiss-dominated with minor metaclastic andrare metavolcanic rocks (Fig. 2c). The Chuacus–Rabinal boundary is a Cretaceous shear zone butthis may represent reactivation of an older structure(e.g. Ortega-Gutierrez et al. 2007). Both complexeshave inherited Proterozoic crust (U–Pb zircon).

Ortega-Obregon et al. (2008) suggested intrusionof the Rabinal granitoids between 462 and453 Ma. Combining all available U–Pb zirconages from the Maya Mountains (Belize), the AltosCuchumatanes (western Guatemala), the ChiapasMassif (southeastern Mexico) and the Rabinalcomplex (see above; Figs 1c & 16a), we suggestthat the Maya Block experienced prolongedOrdovician–Early Devonian magmatism. Weconsider the c. 396 Ma age of a Rabinal orthogneissreliable, as 404–391 Ma magmatites (both graniteand rhyolite) occur north of the Rabinal complex.Our Rabinal augen–gneiss sample is dominatedby 485–400 Ma zircons and shows two age groupsat 465 and 413 Ma. Similar ages occur in theMaya Mountains (c. 418 Ma; Steiner & Walker1996) and in other Rabinal granitoids (c. 462,c. 417 Ma; Ortega-Obregon et al. 2008). TheChiapas Massif and the Rabinal area of centralGuatemala may also host c. 482 Ma granitoids(Weber et al. 2008; Ortega-Obregon et al. 2008;this study).

The Chuacus-complex rocks apparently recyclethese Early Palaeozoic magmatites in their sedimentsand granitoids: Solari et al. (2009) reported, amongothers, 480–402 Ma (main age group at 436 Ma)detrital zircons in a Chuacus-complex paragneiss,and we found c. 403 Ma, probably metamorphiczircon in Triassic orthogneiss. The Chuacus com-plex shows high-P, perhaps ultrahigh-P (Ortega-Gutierrez et al. 2004) metamorphism that is, atbest, constrained in age to between c. 440 Ma, theyoungest significant zircon-age group of theChuacus metasediments and possibly 302 Ma, theage of post-high-P migmatite leucosome, andconservatively to between c. 440 (see above) and238–218 Ma, the age of Triassic orthogneiss(Figs 2d & 16a; see above). The Rabinal complexapparently lacks high-P metamorphism but this hasto be confirmed by quantitative petrology. LowerCarboniferous–Permian Santa Rosa Group sedi-ments cover the Rabinal complex unconformably(Fig. 2c).

Recent studies within and along the southernmargin of the Maya Block suggest a subdivisionof its lithostratigraphy into pre-Early Ordovicianclastic rocks (e.g. Jocote complex, intruded byc. 482 Ma granite in the southern Chiapas Massif,Weber et al. 2008; post-c. 520 Ma Baldy unit ofthe Maya mountains, Martens et al. 2010), and ayounger, again mostly clastic sequence that hasrhyolite intercalations (c. 404 Ma; Bladen For-mation of the Maya mountains, Martens et al.2010) and is intruded by granitoids as young asc. 404 Ma (Maya mountains, Steiner & Walker1996). Ortega-Obregon et al. (2008) suggested thepresence of pre-Middle Ordovician basement inthe Rabinal area of central Guatemala (San


Gabriel unit intruded by the Rabinal granitoidsat c. 462 Ma); Solari et al. (2009) extended theSan Gabriel unit to western Guatemala. Our infer-ence that intrusions in the Rabinal area span482–396 Ma suggests similar prolonged volcano-clastic deposition and intrusion for the Rabinalcomplex and firmly ties this complex to theMaya Block. All areas of the Maya Block exposingOrdovician–Early Devonian granitoids are cove-red by the Santa Rosa Group; characteristicallythese sediments contain Early Palaeozoic detritalzircons. Tectono-stratigraphically, the Rabinal com-plex probably represents continental crust over-printed by an arc to back-arc setting (see alsogeochemical data reviewed in Ortega-Obregonet al. 2008). Detrital zircon geochronology ties theMaya Block to the South American side of theIapetus–Rheic oceans (e.g. Weber et al. 2008 andreferences therein). The Chuacus complex may rep-resent a marginal basin that received detritus fromthe Ordovician–Silurian arc, and Pan-African(c. 0.5–0.7 Ga), and Grenvillian (c. 0.9–1.2 Ga)

hinterland and was subducted into the mantlelikely in the Devonian.

Next, we discuss possible correlations of theRabinal and Chuacus complexes with the Acatlancomplex of the Mixteca terrane of southernMexico (Fig. 1a & c). The Acatlan complex com-prises a nappe stack containing Middle Proterozoicto Palaeozoic units imbricated during the Palaeo-zoic. To date, seven nappes containing distinctivestratigraphy have been discriminated (Fig. 2i;Yanez et al. 1991; Ortega-Gutierrez et al. 1999;Talavera-Mendoza et al. 2005; Keppie et al. 2006;Vega-Granillo et al. 2007). The stack shares ac. 325 Ma regional metamorphism. The hangingEsperanza suite is dominated by Early Silurian(442 + 5 to 440+14 Ma) mega-crystic, K-feldsparaugen gneiss and contains metasediments andmetabasalts transformed to eclogite. The entiresuite records peak PT at 730–830 8C and 1.5–1.7 GPa attained at c. 430–420 Ma. Partial over-printing occurred at 640–690 8C and 1.0–1.4 GPaprior to c. 375 Ma. The Tecolapa suite consists of

Fig. 16. Tectonic–palaeogeographic evolution models for the North American–Caribbean Plate boundary andPalaeozoic–Cenozoic correlations between southern Mexico and the Chortıs Block. (a) Overview of our new andpublished geochronology from the Maya Block, the Chortıs Block, and southern Mexico. Data are referenced in thecaptions of Fig. 1 and in the text.


Fig. 16. (Continued) (b) Permo-Triassic ages of southern Mexico and northern Central America. Ages in italics refer todetrital zircons. (c) Palaeogeographic model for the Permian modified from Weber et al. (2006) and Vega-Granillo et al.(2007) emphasizing arc magmatism along the margins of Laurentia and Gondwana and a common position of theAcatlan complex and the Maya Block in the inner orogenic zone of the Acadian orogen. (d) Jurassic ages along thenorthern Caribbean Plate boundary. Subdivision of the Chortıs Block follows Rogers et al. (2007c). Mineralabbreviations as in Fig. 1. (e) Middle to Late Jurassic tectonics of the Caribbean emphasizing rifting of the Chortıs Blockfrom southern Mexico and its connection with the Arteaga complex of the Zihuatanejo terrane of southwestern Mexico.


Middle Proterozoic mega-crystic granitoids andtonalitic gneisses and Early Ordovician leucogra-nites (478–471 Ma). The greenschist-facies ElRodeo suite contains Cambrian–Early Ordovicianclastics and basalts of continental-rift affinity andis intruded by Ordovician (476–471 Ma), Devonian(371 + 34 Ma), and Early Permian (287 + 2 Ma)granitoids. The Xayacatlan suite comprises eclogite-facies, mafic and pelitic rocks and is intruded byOrdovician leucogranites (478+5 to 461 + 9 Ma),post-dating high-P metamorphism. Peak PT con-ditions were at 490–610 8C and 1.2–1.3 GPa,reached during Early Ordovician (c. 490–477 Ma).The suite re-equilibrated in greenschist faciesduring the Mississippian (c. 332–318 Ma). TheIxcamilpa suite consists of greenschists and blue-schists (200–580 8C, 0.65–0.9 GPa) intercalatedwith metaclastic rocks that are younger thanc. 470 Ma. The high-P event is loosely constrainedbetween 450 and 330 Ma; the absence of Silurian–Devonian detrital zircons and the existence of amajor magmatic event of Silurian age withinseveral units of the Acatlan complex, which maybe genetically linked to subduction, suggest thathigh-P metamorphism was Late Ordovician–EarlyDevonian. The greenschist-facies Cosoltepec suite

is the most widespread unit of the Acatlan complexand consists of monotonous, greenschist-faciesclastic rocks that are younger than c. 410 Maand basalts. This passive margin sequence wasimbricated with the other nappes before the EarlyCarboniferous (Figs 2i & 16a, Acatlan).

We see several gross similarities between theAcatlan-complex suites and the Rabinal andChuacus complexes (Fig. 2c, d & i): (1) the uppernappe stack, from the Xayacatlan suite at thebottom to Esperanza suite at the top, contains exclu-sively Proterozoic detrital zircons, the lower stackalso Ordovician and younger zircons (Magdalena,�275 Ma; Cosoltepec, �410 Ma; Ixcamilpa,�477 Ma). The Maya-Block rocks (Baldy unit,Maya Mountains; Jocote complex, southeasternChiapas), including the Rabinal complex (SanGabriel unit, Guatemala), contain exclusively Pro-terozoic detrital zircons as well, the Chuacuscomplex has a major component of Early Ordovi-cian–Early Devonian detrital zircons (480–402 Ma; c. 436 Ma) besides Proterozoic ages. (2)Although elucidating the Proterozoic history ofnorthern Central America is beyond the topic ofthis paper, we note that the major Proterozoic detri-tal zircon age clusters in both the upper stack suites

Fig. 16. (Continued) (f) Early Cretaceous ages of southern Mexico and northern Central America. (g) Late EarlyCretaceous tectonics of the Caribbean underlining emplacement of the South El Tambor complex rocks onto the ChortısBlock and the existence of continuous arcs off southwestern and southern Mexico: Chortıs–Xolapa–Taxco–Viejo;South El Tambor complex–Teloloapan; Arteaga (Zihuatanejo)–Sanarate (Chortıs).


of the Acatlan complex and the Maya Block/Rabinal complex are Grenvillian (c. 0.95–1.25 Ga) and South American (c. 1.3–1.8 Ga;Rondonia–San Ignacio and Rıo Negro–Juruenaprovinces; see Weber et al. 2008 for locations;Fig. 2c & i). (3) All upper stack suites (Xayacatlanto Esperanza) have massive Early Ordovician to

Middle Devonian intrusions, whereas the lowerstack suites (Cosoltepec to Ixcamilpa) received theerosional detritus of these granitoids. Similarly,the Maya Block/Rabinal complex carries EarlyPalaeozoic magmatism, whereas the Chuacuscomplex collected its detritus. Items (1)–(3) linkthe Maya Block/Rabinal complex to the Xayacatlan

Fig. 16. (Continued) (h) Block diagrams illustrating north–south swath across the Motagua suture zone at c. 70 Ma.Left: the southern Maya Block is imbricated by sinistral transpression. Centre: truncation of the continental margin bysinistral transpression. Roll-back and break-off of the North El Tambor complex slab may have caused magmatism inthe southern Maya Block. Right: hanging wall arc magmatism terminated contemporaneously to slightly after(reflecting transfer to the Bahamas continental margin) their collision with the southern Maya and southern ChortısBlocks (see text for further discussion). Hanging wall arc diagram modified from the Cordillera Central, Hispaniola,situation (Escuder-Viruete et al. 2006a). (i) Geochronology of the footwall and hanging wall of the Motagua suturezone. Left: ages of imbrication and associated metamorphism in the southern Maya Block (this paper). Centre: ages inthe North El Tambor complex (see references in the text), and the subduction–accretion complexes in Cuba(Garcia-Casco et al. 2002, 2007; Schneider et al. 2004; Krebs et al. 2007). Right: ages from the Greater Antilles arc,Cuba, and Cordillera Central arc, Hispaniola (Grafe et al. 2001; Hall et al. 2004; Stanek et al. 2005; Kesler et al. 2005;Rojas-Agramonte et al. 2006; Escuder Viruete et al. 2006b). (j) Subduction–accretion and arc complexes of Cuba andHispaniola and major age groups emphasizing the evolution of the hanging wall of the Motagua suture zone. (k) LateCretaceous tectonics of the northern Caribbean highlighting the two Late Cretaceous suture zones on the Chortıs Block.


Fig. 16. (Continued) (l) Late Cretaceous and Early Cenozoic ages along the northern Caribbean Plate boundarybringing the Chortıs Block to its pre-c. 40 Ma position off southern Mexico. Geochronological data compiled fromHerrmann et al. (1994), Moran-Zenteno et al. (1999), Ducea et al. (2004a), Cerca et al. (2007) and references therein.Deformation zones (numbers 1–9) and their ages are discussed in the text. Convergence direction at c. 45 Ma is fromPindell et al. (1988). (m) Middle Miocene ages along the northern Caribbean Plate boundary. Geochronology is fromthis paper, Moran-Zenteno et al. (1999), and Jordan et al. (2007) and references therein. Deformation zones (numbers1–5) and their ages are discussed in the text.


to Esperanza stack, and the Chuacus complex to theCosoltepec and Ixcamilpa stack of the Acatlancomplex. In both areas, southern Mexico and Guate-mala, Permian, Triassic and Jurassic magmatismstitches the various units together (Figs 1c, 2c, d &i). In contrast, eclogite-facies metamorphism is con-fined to the upper stack in the Acatlan complex, andto the Chuacus complex in Guatemala; the lowerstack in Mexico shows blueschist-facies meta-morphism only. As outlined above, upper stackhigh-P metamorphism is Ordovician (Xayacatlansuite) and Silurian (Esperanza suite), and thelower stack (blueschist-facies) high-P metamorph-ism is 450–330 Ma (best guess of Vega-Granilloet al. 2007 is 458–420 Ma). Our loose constrainfor the Chuacus complex eclogite-facies event is440–302 Ma (see above). These age constrains tiethe Chuacus complex high-P event either to theEsperanza or the Ixcamilpa suite; disparate litho-stratigraphy and age of metaclastic sequence makethe Esperanza suite–Chuacus complex correlationunlikely (Fig. 2c, d & i).

Can we go further with the correlations? (1) Themost obvious correlation, based on lithostratigraphyand type and age of magmatism, compares theRabinal complex with the Esperanza suite of theAcatlan complex. In this scenario, at least the pre-Silurian parts of the Rabinal complex shouldcontain high-P relicts. However, placing theRabinal complex (and the Maya Block) furtherinto the hinterland of the Ordovician orogeny thanthe Esperanza suite would allow variable meta-morphic overprint. (2) We predict that theChuacus complex will see separation into severalunits. For example, several samples of Chuacus-complex orthogneiss (data not shown in thispaper) yielded Proterozoic ages (0.98–1.2 Ga) thatare not or are little rejuvenated by Palaeozoic–Mesozoic events; it is possible that these gneissesdefine a unit that corresponds to the Tecolapa suiteof the Acatlan complex. If our correlations arecorrect, where are the El Rodeo (?arc and continen-tal rift) and Xayacatlan (oceanic basin) suites ofsouthern Mexico in Guatemala? First, these suitesmay represent relative small-scale, regionally con-fined successions (as several units of the CretaceousMixteca–terrane arc succession; see below).Second, only few parts of the Chuacus complexand the Maya Block/Rabinal complex have beenstudied in detail. The abundance of mafic rocks inthe some parts of Chuacus complex and the strongvariability of lithology within it (compare, forexample, the amphibolite–paragneiss-dominatedEl Chol and orthogneiss and migmatite-dominatedRıo Hondo areas described in the Structure andKinematics section) suggests that other suites,reminiscent to those of the Acatlan complex, maybe discovered. Here we propose that the Chuacus

complex that is rich in mafic rocks best correlateswith the Ixcamilpa suite. These units formed a mar-ginal basin that was transferred into a Silurian–Carboniferous subduction–accretion complex; theChuacus complex may have represented the lead-ing edge of this subduction complex. The clastics-dominated Chuacus complex may correspondto the Cosoltepec suite of the Acatlan complex.(3) Although data from the Chortıs Block are stillsparse, our U–Pb analysis identified c. 400 Ma crys-tallization ages in the basement rocks, suggestingthat this block also experienced Early Devonianmagmatism, identical to the youngest ages fromthe Maya Block/Rabinal complex.

We propose that the southern Maya Block/Rabinal complex and the Chuacus complex wereparts of the Taconic–Acadian orogen. The Chua-cus complex due to its high-P metamorphism andRabinal complex due to its extensive Ordovician–Silurian magmatism were located together withthe Acatlan complex within the orogenic interioralong the northern margin of South America. It islikely that the Chortıs Block also belonged to thisorogen (Fig. 16c). Subduction produced an EarlyOrdovician arc magmatism on the Tecolapa and ElRodeo suites, which was followed by high-P meta-morphism and accretion of the Xayacatlan suiteto the Tecolapa and El Rodeo suites. Subductioncontinues beneath the Xayacatlan producing thepost-high pressure granitoids within the Xayacatlansuite. Subsequently the Maya Block/Rabinal com-plex and the Esperanza suite were accreted andthe Esperanza suite involved in continental sub-duction. Subduction again migrated outward, pro-ducing the Ixcamilpa and Chuacus complexes andinvolved them in subduction. In the CarboniferousLaurentia and Gondwana collided producing thenappe stack of the Acatlan complex and Chuacuscomplex.

Permo-Triassic arc on Maya and

Chortıs Blocks

Gneiss of the Rabinal complex of westernGuatemala was migmatized at c. 218 Ma and inher-ited c. 268 Ma zircon; thus the high-grade rocksthat accompany the low-grade metasedimentaryrocks of the San Gabriel unit are probably a resultof local Triassic (and Jurassic magmatism, seebelow) migmatization. In central Guatemala,Chuacus-complex orthogneiss, intruding high-Pparagneiss and garnet amphibolite (relict eclogite),records two zircon-crystallization events at 238and 219 Ma. Regional magmatism and metamorph-ism is demonstrated by similar hornblende andwhite mica cooling ages (238–212 Ma; Fig. 16a &b). Phyllite interlayered with conglomerate attribu-ted to the Santa Rosa Group of central Guatemala


(Rabinal–Salama area) contains detrital zircons atc. 282, c. 260 and c. 245 Ma. Volcanic pebbles dom-inate the conglomerate. Triassic deformation in theRabinal and Chuacus complex shows high-T flowwith north(east)–south(west) shortening and bothsub-vertical and along-strike, NW–SE extension.In the Chortıs Block, migmatitic orthogneiss crys-tallized at c. 273 Ma and underwent high-T meta-morphism at c. 245 Ma. Inherited c. 241 Mazircon in Jurassic orthogneiss and biotite coolingages support regional Triassic magmatism andmetamorphism (Fig. 16a & b).

Permo-Triassic granitoids are documented alongthe length of Mexico and probably extend intonorthwestern South America (e.g. Centeno-Garcıaet al. 1993). Their apparent linear arrangement prob-ably traces an active continental margin that origi-nated after final amalgamation of Pangaea alongthe Ouachita–Marathon suture at c. 290 Ma(Fig. 16c; Torres et al. 1999). Isotopic ages ofthese granitoids span 287–210 Ma with reliableU–Pb zircon ages at 287–270 Ma (Figs 1c & 16a,Ttoletec/Magdalena/Oaxaca, 16b; Yanez et al.1991; Solari et al. 2001; Elıas-Herrera & Ortega-Gutierrez 2002; Ducea et al. 2004a). In contrast tothese mostly undeformed granitoids, the ChiapasMassif batholith rocks (Fig. 16b) are deformed andmetamorphosed (Weber et al. 2006). Rock typesinclude orthogneisses that are intruded by calc-alkaline rocks varying from granite to gabbro,metasedimentary rocks (metapelite/psammite, calc-silicate; Sepultura unit) and mafic rocks (hornblendegneiss; Custepec unit). Zircons from orthogneissyielded a concordia age at 271.9 + 2.7 Ma, andthree samples from migmatitic paragneiss and leuco-some yielded 254.0 + 2.3 to 251.8 + 3.8 Ma ages(Fig. 16a, Chiapas). The latter ages date the mostprominent event in the Chiapas Massif – partialanatexis in pelite/psammite and, locally, ortho-gneiss. High-grade metamorphism is contempora-neous with c. north–south shortening (Weber et al.2006).

Our new ages indicate that during the Permo-Triassic the Rabinal and Chuacus complexes andthe northern Chortıs Blocks were part of a conti-nuous magmatic arc from northeastern Mexico tonorthern South America (e.g. Dickinson & Lawton2001), which initiated after the formation ofPangaea along the western continental margin ofGondwana (Fig. 16c). Sample 5H-9a zircons of mig-matitic orthogneiss in the north-central ChortısBlock have within error the same intrusion and high-grade metamorphic ages as the Chiapas Massif rocks(Fig. 16a). This identical geological history suggestsa direct correlation between the Chiapas Massif andthe northern Chortıs Block along the Permo-Triassicarc prior to the opening of the Gulf of Mexico(Fig. 16b & c). The ages in the Rabinal and

Chuacus complexes are c. 30 Ma younger thanthose in the Chiapas Massif and also younger thanmost Permo-Triassic ages in southern Mexico. Wespeculate that the Chuacus complex occupied,together with the Maya Mountains of Belize (c.230 Ma reheating), an easterly position along themain Chiapas–Chortıs arc (Fig. 16c). The youngerages may reflect an episode of flat-slab subductionthat shortened the main arc (Weber et al. 2006)and caused a younger episode of magmatism andmetamorphism in its hinterland. Shorteningdirections are similar in the Rabinal and Chuacuscomplexes and the Chiapas Massif. Associatedwith the shortening there is probably a dextral-transpressional component, widespread in southernMexico (e.g. Malone et al. 2002), which is consist-ent with tangential stretching in the TriassicChuacus complex. Finally, the conglomerate claststhat are predominantly sourced from intermediatevolcanic rocks, and c. 245, c. 260 and c. 280 Ma det-rital zircons within phyllite demonstrate that parts ofthe clastic cover sequence of the Rabinal complex incentral Guatemala is not Lower Carboniferous–Permian Santa Rosa Group. We suggest that a partof this Group comprises either a new unit or ispart of the undated basal conglomeratic sequenceof the Jurassic–Cretaceous Todos Santos Formation(Fig. 2b & d). Detritus and age of detrital zirconsmake the Chiapas–Chortıs-arc Massif the likelysource of these molasse-type, basal deposits.

Jurassic rifting and arc magmatism

Granite intruded into the Rabinal complex ofwestern Guatemala at c. 171 Ma (164–177 Mazircons; Fig. 16a). Most of the Maya Block wasprobably emergent during the Triassic–Jurassicwith the likely Middle Jurassic–Early CretaceousTodos Santos Formation reflecting rifting duringJurassic separation of North and South America(Fig. 2b; e.g. Donnelly et al. 1990). Granites alsointruded the northern Chortıs Block during theMiddle Jurassic (c. 159 and 166 Ma, zircons at140–191 Ma; 168 Ma; Fig. 16d). The Middle Juras-sic, clastic, partly greenschist-facies Agua FrıaFormation (Fig. 2j) thickens from the central tothe eastern Chortıs Block and is interpreted to berelated to transtension and rifting during the Jurassicopening between North and South America; e.g.the Guayape fault zone may have originated as aJurassic normal fault (Fig. 16d; e.g. James 2006;Gordon 1993). The rocks of the Agua Frıa For-mation were deformed during the Late Jurassic(Viland et al. 1996).

Early–Middle Jurassic magmatism–metamor-phism affected the Chazumba and Cosoltepecunits of the Acatlan complex and produced the Mag-dalena migmatites of southern Mexico (Figs 2i &


16a, Totoletec/Magdalena/Oaxaca; e.g. Talavera-Mendoza et al. 2005; our new c. 162 Ma coolingage); inferred metamorphic conditions were low-P/high-T (c. 730 8C at c. 0.55 GPa; Keppie et al.2004). Anatectic granitic dykes of the San Miguelunit (175–172 Ma; e.g. Yanez et al. 1991) cut theformer units. These events have been interpreted asthe result of a plume breaking up Pangaea, openingof the Gulf of Mexico and Triassic–Middle Jurassicmagmatic arc activity (e.g. Keppie et al. 2004).Approximately at 165 Ma, NNW-trending, dextraltranstension shear, kinematically compatible withopening of the Gulf of Mexico, occurred alongthe Oaxaca fault (Alaniz-Alvarez et al. 1996), andNW-trending, Jurassic–earliest Cretaceous dextraltranspression took place within the northernOaxaca complex (Solari et al. 2004). The Xolapacomplex (Fig. 2l) also shows 165–158 Mamagmatic–metamorphic ages (Fig. 16a, Xolapa;Herrmann et al. 1994; Ducea et al. 2004a). TheZihuatanejo and Huetamo mature arc sections ofthe Guerrero terrane rest unconformably on theArteaga–Las Ollas basement, which comprisesocean-floor–continental-slope–subduction com-plexes that were deposited in the Late Triassic–Early Jurassic, intruded by 187 + 7 Ma granite,and deformed in the Middle Jurassic (Fig. 2m).Major late Middle–Late Proterozoic–Phanerozoicdetrital zircon populations from these complexesare c. 247, c. 375, c. 475, c. 575 and c. 975 Ma. Theoverlying Cretaceous arc has major zircon popu-lations at c. 160, c. 258, c. 578 and c. 983 Ma(Fig. 2m; Talavera-Mendoza et al. 2007).

Our tectonic model (Fig. 16e) integrating theJurassic evolution of the Maya and Chortıs Blockswith the magmatic–tectonic evolution of southernMexico is modified from Mann (2007). Onset ofrifting in the Gulf of Mexico and between Northand South America reduced Triassic (see above) toJurassic volcanic eruption rate in southern Mexicoand the Maya and the Chortıs Blocks. Late intrusioninto active dextral fault zones (Mexico) may explainthe concentration of magmatism at 175–166 Ma(Fig. 16d). Rifting of the Chortıs Block fromsouthern Mexico probably initiated along thearc and developed either as back-arc extensionabove the proto-Pacific subduction zone, a failedarm of the Proto-Caribbean rift, or a combination.Additionally, the Arteaga–Las Ollas basementunits rifted from southern Mexico, as they likelyshare the Late Triassic–Jurassic history of southernMexico (Fig. 2m; e.g. Talavera-Mendoza et al.2007). Our model requires rifting the ChortısBlock from southern Mexico and creation ofoceanic crust, because MORB oceanic crust andmelange of the South El Tambor complex ispresent between the Chortıs Block and southernMexico. This must predate 132 Ma, the age of

eclogite formation in the South El Tamborcomplex and is probably pre-Late Jurassic, the prob-able age of the oldest radiolarians in the South ElTambor cherts (Chiari et al. 2006). Rifting affectedalso the interior of the Chortıs Block (Agua Frıa For-mation; Fig. 16d). These rift structures trend c. NE(present coordinates), sub-parallel to the northwes-tern and southeastern margins of the Chortıs Blockand the Guayape fault. This rifting is distinctlydifferent from the WNW-trending Mid-Cretaceousintra-arc rifting (Rogers et al. 2007a; see below).The Late Jurassic regional metamorphism (165–150 Ma; Fig. 16a) and deformation (Viland et al.1996) may be rift and strike–slip related.

Several likely connections exist between theChortıs Block, the Xolapa complex, and the Guer-rero composite terrane. The Sanarate complex(Fig. 2e), the westernmost basement outcrop of theChortıs Block south of the MSZ, was includedwith the South El Tambor complex because of itsmetamorphism (lower metamorphic grade thanthat of the Las Ovejas unit) and rock assemblage(dominated by mafic and clastic rocks; e.g.Tsujimori et al. 2004, 2006); however, the Jurassicage for the metamorphism precludes such a corre-lation. Here we suggest a correlation of this com-plex with the Arteaga–Las Ollas units (Figs 2m &16e) based on lithology, likely proximity to theJurassic continental margin of Mexico and distinctJurassic metamorphism (Fig. 2e & m). The Sanaratecomplex probably also includes parts of the Cretac-eous Zihuatanejo and Huetamo volcanic and clasticunits (Fig. 2m). Both the Arteaga–Las Ollas unitsand the overlying Cretaceous arc contain zirconsthat may have originated from the Chortıs Block.The c. 247, c. 257 and 475 Ma clusters, particularlyin the �85 Ma arc rocks, are difficult to derivefrom North America because of intervening spread-ing ridges and subduction zones (Figs 2m & 16e, seebelow). The pre-migmatitic basement of the Xolapacomplex, containing evidence for a Grenvillecrustal component (Herrmann et al. 1994) andPermian magmatism (see above) in an alternationof metagreywacke and metabasic rocks, probablycorrelate to the Permian and Jurassic section of theMixteca terrane (Fig. 2m, Olinala unit). Finally,connecting the Jurassic magmatic–metamorphicrocks of southern Mexico and the northern ChortısBlock implies �1000 km post-Jurassic offset.

Early Cretaceous subduction, top-to-south

emplacement of the South El Tambor complex,

and proto-Pacific arc formation

Spreading between the Chortıs Block and southernMexico must have reverted to subduction before132 Ma (see above). The c. 120 Ma phengite ages


of the South El Tambor complex blueschists areinterpreted as marking fluid infiltration during retro-gression from eclogite-facies conditions (see above)and could date exhumation during emplacement ofthe South El Tambor complex melange onto theChortıs Block (Fig. 8b). The El Tambor rocks areemplaced top-to-south (present coordinates).Aptian–Albian calc-alkaline, arc volcanic rocks ofthe Manto Formation and volcanoclastic rocks ofthe Tayaco Formation, sandwiched betweenshallow water carbonates, were deposited in azone of intra-arc extension in the central ChortısBlock, the Agua Blanca rift (Rogers et al. 2007a;Figs 2j & 16f). The strike of the rift is WNW(present coordinates), oblique to the northernmargin of the Chortıs Block and the Jurassic exten-sional structures (see above). Cretaceous (peaks at c.128, c. 118 and c. 81 Ma; Figs 16a, f & 5h) mag-matic rocks are widespread on the Chortıs Block(Fig. 16f) and probably mark continuous arc for-mation above the proto-Pacific subduction zone(Fig. 16g). The Xolapa complex, the largest plutonicand metamorphic mid-crustal basement unit insouthern Mexico, preserves abundant migmatitesthat were produced during a single metamorphicevent with peak prograde conditions at 830–900 8C and 0.63–0.95 GPa. This occurred within acontinental magmatic arc and overprinted a pre-migmatitic metamorphic assemblage (Corona-Chavez et al. 2006). Deformation, likely (north-)northeast–(south-)southwest shortening (Corona-Chavez et al. 2006), associated with low-Pmetamorphism and migmatization predates 129 +0.5 Ma in the Tierra Colorada area of the westernXolapa complex (Solari et al. 2007). The migma-tites have been directly dated at 131.8 + 2.2 Ma(Herrmann et al. 1994) and 130–140 Ma (Duceaet al. 2004a). The structural top of the migmatizedsequences is the 126–133 Ma Chapolapa Formation(Campa & Iriondo 2004; Hernandez-Trevino et al.2004) that mostly comprises andesite/rhyolite; itshanging wall, Albian–Cenomanian Morelos For-mation shelf limestone, is again in tectonic contact(Figs 2l & 16a, Xolapa; e.g. Riller et al. 1992;Solari et al. 2007).

Rogers et al. (2007a) suggested several featuresand piercing lines that tie the Chortıs Block tosouthern Mexico; here we elaborate on the tectonicmodel explaining the correlation (Fig. 16g). Wesuggest that the 141–126 Ma migmatization–magmatism of the Xolapa complex represents arcmagmatism related to subduction of the South ElTambor complex lithosphere. Characteristically,no Cretaceous ages younger than c. 126 Ma havebeen found in Xolapa complex, so the end of arcmagmatism in the South El Tambor subductionzone is likely to be contemporaneous with emplace-ment of the South El Tambor Complex onto the

Chortıs Block (�120 Ma, see above). Laterally,Xolapa-complex magmatism is correlated withthat in the Taxco–Viejo unit of the Mixtecaterrane, also an arc on continental crust (e.g.Centeno-Garcıa et al. 1993) and dated at130.0 + 2.6 and 131.0 + 0.85 Ma (Campa &Iridono 2004). This correlation is supported bythe, within error, identical ages of the ChapolapaFormation meta-andesite/rhyolite (structural topof the Xolapa complex) and the volcanic flows ofthe Taxco–Viejo unit (Fig. 2l & m). Magmatismin the Teloloapan intra-oceanic arc ended at c.120 Ma. Magmatism in the oceanic arc andback-arc region of the Arcelia units, outboard(southwesterly) of the Teloloapan unit continuedinto the Cenomanian (Fig. 2m). The youngest detri-tal zircon fractions in the clastic rocks above the vol-canic succession in the Teloloapan and Taxco–Viejo units, comprising magmatic zircons fromlocal sources, are c. 124 and c. 132 Ma, respectively(Talavera-Mendoza et al. 2007), thus contempora-neous with the Xolapa-complex zircons. Again,these ages indicate that magmatism terminated atthat time and suggest a close spatial relationshipof the Taxco–Viejo–Xolapa–Teloloapan arcs(Figs 2m & 16g). The NE-dipping subductionzone of the South El Tambor complex may haveended at a transform fault at the eastern end of theMaya Block (Fig. 16g).

The Teloloapan, Taxco–Viejo, and Olinala unitsand the Xolapa complex are covered by Albian–Lower Cenomanian Morelos shelf limestones(Fig. 2l & m). These limestones may correspondto the massive La Virgen limestone atop theSouth El Tambor complex (Fig. 2k). Carbonatepassive margin strata (Coban–Campur Formations;Fig. 2b) were deposited on the Maya Block. Thepost-suturing deposits on the Chortıs Block areprobably the red beds and foreland-basin depositsof the Valle de Angeles and Gualaco Formations.The latter is typically intercalated with UpperCretaceous volcanic rocks indicating establishmentof the Sierra Madre Occidental arc in Late Cretac-eous, also on the Chortıs Block (see below). Charac-teristically, proto-Pacific subduction continuedbeneath the Chortıs Block and the Zihuatanejounit far into the Late Cretaceous (Fig. 2j & m). Inour interpretation the Agua Blanca rift and theManto volcanic rock intercalations are related toproto-Pacific subduction and intra-arc extension.We attribute the inversion of the Agua Blancaintra-arc rift to the accretion of the Zihuatanejo–Huetamo units–Sanarate complex to southernMexico and the western Chortıs Block. Top-to-eastshear in the greenschist facies, metabasaltic/clasticSanarate complex of the western Chortıs Blockreflects this accretion of the Guerrero terrane/Sanarate complex.


Why were so many individual arcs and subduc-tion–accretion complexes developed along thesouthwestern margin of Mexico (the Guerrero com-posite terrane) and the Chortıs Block (the Sanaratecomplex) during the Early and Late Cretaceous(e.g. Talavera-Mendoza et al. 2007)? If riftingbetween the Chortıs Block–Arteaga unit andsouthern Mexico was mostly related to spreadingbetween the two Americas and formation of theGulf of Mexico (Fig. 16e), spreading was probablynorth–south with c. north-trending transformfaults. Transformation of this rift to convergencealong the c. NNW-trending proto-Pacific subduc-tion zone may have allowed activation of sometransform faults as subduction zones.

Late Cretaceous North El Tambor complex

subduction and Maya Block–Caribbean

arc collision

Our structural, petrological, and geochronologicaldata demonstrate sinistral transpressive collisionalong the southern margin of the Maya Blockduring Late Cretaceous. Collision transferred thesouthern Maya Block into a nappe stack with theNorth El Tambor ophiolitic melange at top. Charac-teristically, deformation involved north–southshortening and east–west lengthening. Sinistraltranspression was partitioned on all scales betweenoblique stretching along the foliation and sinistralstrike–slip along steeply dipping shear zones(Fig. 16h). To a first order, deformation is identicalin the North El Tambor complex, Chuacus complex,and the northern foreland fold–thrust belt; 450–550 8C and 0.7–0.8 GPa regional Cretaceousreheating within the Chuacus complex, associatedwith local magmatism, is synkinematic. Penetrationof deformation and overburden (PT conditions)increases from west to east along the collisionzone. The relatively high-P/low-T metamorphismand nearly coeval blueschist- (North El Tamborcomplex) and amphibolite-facies metamorphism(Chuacus complex) suggest a cold geotherm andsteep continental subduction (Fig. 16h). SouthernMaya-Block imbrication is dated at 75–65 Ma(Fig. 16i) and cooling to upper crustal levels(c. 100 8C) was attained by Early Cenozoic (AFTresults at high elevation, see above). The Jurassic–Early Cretaceous North El Tambor oceanic andimmature arc rocks incorporated c. 131 Ma eclogiteinto an accretionary wedge that reached crustallevels by Late Cretaceous (c. 76 Ma).

What collided with the southern Maya Block inthe latest Cretaceous? We follow and broaden themodel of Mann (2007) and Rogers et al. (2007b)and suggest that the Maya Block–North ElTambor complex suture extends toward the west

into the footwall of the Siuna terrane of the southernChortıs Block and the Nicaragua Rise and to the eastinto the subduction–accretion and arc complexes ofJamaica, Hispaniola (Dominican Republic) andCuba (Fig. 16j & k). In comparison with the NorthEl Tambor complex (Fig. 2a), the Siuna melange(Fig. 2f) contains similar lithology (ultramaficrocks, radiolarite, meta-andesite, volcanic arenites)of proven Late Jurassic age, high-P phengites at c.139 Ma, blueschist and eclogite, and Aptian–Albian turbiditic sequence. These rocks are overlainby the Atima Formation and intruded by diorite (c.60 Ma) and andesite dykes (c. 76 Ma; Fig. 2f;Flores et al. 2007; Venable 1994). The poorlydated (post-80 Ma) Colon fold belt of the easternChortıs Block and the Nicaragua Rise probablyreflects imbrication along the footwall of the Siunaterrane (Rogers et al. 2007b). The Siuna terraneprobably extends westward into the ocean igneouscomplexes of the Santa Elena and Nicoya peninsu-las of Costa Rica; these complexes contain c.165–83 Ma pillow lavas, sills and flows, radiolarianchert, basalt breccias, and gabbro and plagiogranite.They are unconformably capped by �75 Ma reeflimestones (Figs 2g & 16i; e.g. Hoernle & Hauff2005). The Cuban subduction–accretion complexcontains c. 148–65 Ma high-P rocks (major peaksat c. 116 and c. 70 Ma; Fig. 16i & j; Krebs et al.2007, Schneider et al. 2004). The Greater Antillesarc of Cuba was active at 133–66 Ma (Fig. 16i &j; major peaks at c. 72 and c. 80 Ma; Grafe et al.2001; Hall et al. 2004; Rojas-Agramonte et al.2005; Stanek et al. 2005), and the CordilleraCentral arc (Dominican Republic) spannedc. 117–74 Ma (Fig. 16h–j; major peaks at c. 75,c. 84 and c. 116 Ma; Kesler et al. 2005; EscuderViruete et al. 2006b). Thus these subduction–accretion complexes and arcs cover time rangescomparable to that of the North El Tamborcomplex, with subduction and arc formation termi-nated c. 10 Ma after collision along the southernMaya Block. This indicates the time span requiredto transfer these subduction–accretion and arccomplexes to the Bahamas-platform collision zonewhen the southern Maya Block acted as a sinistraltransform fault zone. Accordingly, terminationof pronounced magmatism in the Santa Elena–Nicoya subduction-accretion complexes is approxi-mately 10 Ma earlier (Fig. 16i) and reflects theirearlier collision with the southwestern ChortısBlock. Based on the termination of magmatism at74 Ma within the Cordillera Central arc and itswell-developed sinistral transpressive deformation(Escuder Viruete et al. 2006a, b), we speculatethis arc was the immediate hanging wall of thecollision zone along the southern Maya Block.

What was the origin of the �75 Ma magmatism(pegmatites) in the southern Maya Block? Causes


might include Pacific-arc magmatism, local anatexisdue to crustal thickening, or break-off of the Siunaterrane–North El Tambor complex–Caribbean arcslab. We favour the latter (Fig. 16h), as entry ofthe Nicoya–Siuna–North El Tambor–Caribbeanslab into the Proto-Caribbean sea must have indu-ced its rapid anticlockwise rotation, rapid along-strike length reduction, and thus slab fracturingand roll-back. In contrast, contemporaneous Pacificarc subduction occurred far to the west on theChortıs Block (Fig. 16k) and the thermal overprintdue to crustal stacking was probably insufficientand too early in the Chuacus complex to causeanatexis (see above).

Cerca et al. (2007) outlined a mostly eastvergent, partly sinistral transpressive, north-trending fold-thrust belt along the western Mixtecaterrane, the Guerrero–Morelos platform that origi-nated during the Coniacian (�90 Ma; Cenomanian,�100 Ma, after Talavera-Mendoza et al. 2005) andwas active into the Cenozoic. They related thiseast–west shortening to the collision of the Carib-bean Plate with the Chortıs Block. Rogers et al.(2007b) and our data show, however, that this col-lision is c. north–south and�76 Ma on the southernChortıs and Maya Blocks (see above). Probably,the Laramide fold–thrust belt orocline on theGuerrero–Morelos platform, along the Sierra deJuarez (Cerca et al. 2007), and within the southernMaya Block (this study) is a composite structure,related to accretion of the most outboard Guerrero-terrane units, Pacific subduction and collision ofthe Siuna terrane–North El Tambor complex–Caribbean Plate. We suggest that shortening in theGuerrero–Morelos platform and its continuationinto the Chortıs Block (Agua Blanca rift inversion)is mainly related to the �80 Ma Arteaga–westernChortıs (Sanarate complex) suture and Pacific sub-duction and the Sierra de Juarez–southern MayaBlock shortening due to the �76 Ma CaribbeanPlate collision (Fig. 16k). Structural superpositionsmust have occurred mostly in the Mixteca–Oaxaca basement terrane area.

Late Eocene–Recent eastward displacement

of the Chortıs Block

Most tectonic models place the Chortıs Block oppo-site southern Mexico during Late Cretaceous–EarlyCenozoic (e.g. Mann 2007; see above). Rogers et al.(2007a, b) summarized features and piercing linesthat support this connection. Here we go overadditional evidence and trace the eastward displace-ment of the Chortıs Block to its present position.Approximately 80–45 Ma magmatic rocks of theSierra Madre Occidental crop out along the coastof southwestern Mexico and inland up to c. 998W(Fig. 16l; Moran-Zenteno et al. 1999; Schaaf et al.

1995). This magmatic arc continued into thecentral Chortıs Block. Lining up the northwesttrend of these rocks in Mexico (the arc continuesNW of the area shown in Fig. 16l) and on theChortıs Block yields �1100 km sinistral displace-ment and �408 anticlockwise rotation of theChortıs Block (Figs 1c & 16l). Based on one datapoint from the central Chortıs Block, the palaeo-magnetically determined anticlockwise post-60 Marotation is c. 328 (Gose 1985). During the Oligo-cene, eastward migration of Chortıs induced a north-eastward migration of the trench (Herrmann et al.1994) and the associated Sierra Madre del Sur arcshifted inland and eastward into southeasternMexico. The arc trends ESE (Moran-Zenteno et al.1999), which is particularly evident on the ChortısBlock (Fig. 16l).

When did the Chortıs Block break away fromsouthern Mexico? In the Chortıs Block, high-gradeand high-strain deformation began at �40 Maand we suggested above that there were majordeformation phases at 40–35 Ma and possibly at25–18 Ma. Sinistral transtension in the La Venta–Tierra Colorada shear zone along the northwesternmargin of the Xolapa complex is low-grade atc. 300 8C (low-T quartz ductility; Riller et al.1992) and started at �45 Ma (Solari et al. 2007).Local high-strain flow occurred at c. 35 Ma (seeabove). Post-tectonic intrusions are 34 + 2 Ma(Tierra Colorada granite, Herrmann et al. 1994);however, brittle, sinistral transtension continuedduring cooling of this granite (Riller et al. 1992).The distinctly different exposure level of the con-tinuous Early Cretaceous Taxco–Viejo (west; upto greenschist facies) and Xolapa arc (east; amphi-bolite facies and migmatization) is ascribed tomajor exhumation of the latter. Large-scale sinistraltranstension along the northwestern Xolapa-complex margin certainly contributed to this differ-ence. Robinson et al. (1990) obtained a c. 40 Ma agefor synkinematic tonalite north of Acapulco. Sinis-tral shear occurred along the northern margin ofthe eastern Xolapa complex between 29 and24 Ma (Chacalapa shear zone, point 4 in Fig. 16l;Tolson 2005). Other sinistral transtensive defor-mation zones in southern Mexico outside theXolapa complex are dated to between c. 37 and c.27 Ma (points 1 and 3 in Fig. 16l). Corona-Chavezet al. (2006) estimated that about 70% of theeastern Xolapa complex was affected by discontinu-ous mylonitic–cataclastic, post-migmatitic defor-mation. We conclude that the migration of theChortıs Block initiated �40 Ma and was fullyactive at 40–35 Ma. This is compatible with thechronology of spreading in the Cayman Troughpull-apart basin (see above).

Using (U–Th)/He thermochronology, Duceaet al. (2004b) estimated average exhumation rates


for the Xolapa complex. A western segment atAcapulco yielded ages between 26 and 8.4 Ma anda rate of 0.22 mm/annum; an eastern profile atPuerto Escondido gave 17.7–10.4 Ma ages and0.18 mm/annum. This slow Neogene exhumationcontrasts with petrological data that suggest exhu-mation from �25 km (migmatites; Corona-Chavezet al. 2006) and 20–13 km (granitoids; Moran-Zenteno et al. 1996) since 34–27 Ma, the age ofthe granitoids; therefore, most exhumation musthave taken place before the Miocene. Pooling the(U–Th)/He and AFT data of Ducea et al. (2004b)and excluding outliers (e.g. ages older than cry-stallization and those with very large errors) indi-cate that most samples of the western profilecooled through 70–100 8C or were exhumed fromc. 4 km at c. 25 Ma. Except for three outliers mostsamples of the eastern profile have identical agesover 2 km elevation and thus suggest rapid coolingat c. 15 Ma (Fig. 6d). Using the three plutons forwhich U–Pb zircon ages are available (Herrmannet al. 1994), we obtained exhumation rates ofc. 1.5 mm/annum between 31 and 25 Ma (12.9 kmdepth at 31 Ma and c. 4 km at c. 25 Ma) andc. 1.9 mm/annum between 32 and 25 Ma (17.9 kmdepth at 32 Ma and c. 4 km at c. 25 Ma) for thewestern section, and 1.2 mm/annum (20.4 kmdepth at 29 Ma and c. 4 km at c. 15 Ma) for theeastern profile. For the western section, Herrmannet al. (1994) obtained two c. 47 Ma migmatizationages; this yields c. 0.8 mm/annum between 47 and31 Ma (c. 26 km depth at 47 Ma and 12.9 kmat 31 Ma). Along the northern rim of the ChortısBlock our data indicate an exhumation rate ofc. 1 mm/annum between 36 and 9 Ma (c. 28 kmdepth at 36 Ma and c. 3 km at 9 Ma). These dataindicate that the early stages of Chortıs-Blockmigration were associated with rapid exhumationdue to sinistral transtension and erosion.

The early stage of distributed sinistral displace-ment and erosion that affected a broad zone insouthern Mexico and the northern Chortıs Blockalso must have affected the southern Maya Block,as its eastern extension, the Nicaragua Rise, waslocated off the western edge of the Maya Block(Fig. 16k). We interpret our AFT data from thesouthern Chiapas Massif and the Chuacus com-plex, both clustering at c. 30 Ma (Fig. 6a & c), torecord this distributed transtension–transpression–sinistral migration and erosion phase.

Why did the Chortıs Block break away at�40 Ma? Following Ratschbacher et al. (1991)and Riller et al. (1992), we propose that the contem-poraneous change in several plate tectonic boundaryconditions triggered break away. During Eocene,shear stresses across the arc were high due to rapidconvergence of relatively young lithosphere underan angle of c. 458 to a pre-existing lithospheric

weakness, the South El Tambor suture. Most impor-tantly, sinistral break away started in an area thathad been an arc at least since Late Cretaceous (seeabove) and thus was thermally weakened. This is acommon process in oblique subduction settings.Consequently, the largest volume (Moran-Zentenoet al. 1999) and the concentration in time(c. 33 Ma Chortıs Block, c. 34 Ma Xolapa com-plex; Fig. 16a) of magmatism and metamorphismseems to have been controlled by the NW–SEshear zones associated with transtensional–transpressional tectonics at the boundary betweensouthern Mexico and the Chortıs Block (Herrmannet al. 1994; Fig. 16l).

Early and Middle Miocene (c. 20–10 Ma) volca-nic rocks are distributed around the Isthmus ofTehuantepec (Fig. 16m). In this region, it isevident that volcanism in the inland regions wasactive after cessation of Oligocene plutonism inthe Xolapa complex (Moran-Zenteno et al. 1999).The earliest manifestations of magmatic activityalong the eastern and central Trans-Mexican volca-nic belt are as old as Middle Miocene (c. 17 Ma;Ferrari et al. 1994) and thus this belt may form adirect continuation of the volcanism around theIsthmus of Tehuantepec. To the east coeval magma-tism occurred along the western part of the Polochicfault zone in westernmost Guatemala (Fig. 16m),indicating �100 km offset along the Polochic faultsince then. This is similar to the c. 130 km offsetproposed along this fault between 10 and 3 Ma(Burkart 1983; Burkart et al. 1987). Still furthereast, the c. 16.9–13.4 Ma ignimbrite provincestrikes northwest–southeast across the ChortısBlock (Fig. 16m; e.g. Jordan et al. 2007). Aligningthe southeastern Mexican and Chortıs Block volca-nic provinces yields c. 400 km displacement sincec. 15 Ma and c. 300 km offset along the Motaguafault (Fig. 16m). Several elements trace the displa-cement during this time (Fig. 16m): tectonic andstratigraphic features NW of the Isthmus of Tehuan-tepec indicate sinistral transtension with c. northtrending pull-apart basins coeval with volcanism(Moran-Zenteno et al. 1999); the easternmost partof the Xolapa complex experienced rapid coolingat c. 15 Ma (Fig. 6d); the Chacalapa shear zonewas reactivated at c. 15 Ma; the Tonala and Polochicshear zones were active at 9–8 and 12–5 Ma,respectively; and most of the AFT ages in theChortıs Block closed at c. 12 Ma (all data fromthis study).

Variations in the Late Cenozoic stress field

North–south rifts that disrupt the ignimbrite pro-vince in western Honduras and southeasternGuatemala commenced after c. 10.5 Ma (Gordon& Muehlberger 1994). They are related to plateau


uplift following slab break-off underneath CentralAmerica (Rogers et al. 2002). Our 11.2–7.9 MaAFT data from the western flank of the Sulagraben support the Late Miocene onset of rifting,although extension structures formed as early asc. 28 Ma, contemporaneous with and kinematicallyrelated to sinistral wrenching along the northernmargin of the Chortıs Block. Numerical modellingsuggests that presently the western edge of theChortıs Block is pinned against North America,making the triple junction between the Cocos,North American and Caribbean plates a zone ofdiffuse deformation (Alvarez-Gomez et al. 2008).Our ‘Late Cenozoic deformation’ structural event(Figs 10, 13 & 15), reflecting sinistral slip alongc. east-striking strike–slip faults locally interactingwith c. north-striking normal faults, is distributedacross the entire plate boundary from the northernforeland to south of the Jocotan–Chameleconfault zone. The transition from older dominantstrike–slip to younger, prevailing normal faulting,proposed by Gordon & Muehlberger (1994),is also evident from our data. However, normalfaulting exemplified by the grabens in southernGuatemala and western Honduras (Fig. 15),extends northward across the Motagua fault zoneinto the Cenozoic Subinal Formation (Fig. 13), theNorth El Tambor rocks on the central Chuacuscomplex (Fig. 13), and the Rabinal complex andCenozoic–Quaternary volcanic rocks of westernGuatemala (Fig. 10). This may indicate that todaythe Polochic fault constitutes the major plate bound-ary fault. This is compatible with termination of theactive strike–slip and reverse fault province ofeastern Chiapas north of the Polochic fault zone(Andreani et al. 2008). Furthermore, our datasupport additional salient features of the Late Ceno-zoic stress distribution: shortening directions trendc. NW–SE in western and west-central Guatemalaand in particular in western Guatemala south ofthe Polochic fault (‘Late Cenozoic deformation’;Figs 10, 13 & 15) and outline the dextral, WNW-trending Jalpatagua distributed fault zone alongthe active volcanic arc (Gordon & Muehlberger1994; Alvarez-Gomez et al. 2008).

Conclusions

New geochronological, petrological, and structuraldata constrain the P–T–d– t history of the northernCaribbean Plate boundary in southern Mexico,central Guatemala and northwestern Honduras. Inparticular, we investigated which crustal sectionsof the Motagua suture zone accommodated thelarge-scale sinistral offset along the plate boundary,and the chronological and kinematic framework ofthis process. Correlations between the Maya andChortıs Blocks, tectono-stratigraphic complexes of

southern Mexico and arc/subduction complexesof the northwestern Caribbean constrain modelsfor the Early Palaeozoic to Recent evolution of theNorth American–Caribbean Plate boundary.

During the Early Palaeozoic, southern Mexico(Acatlan complex of the Mixteca terrane), thesouthern Maya Block (Rabinal complex), theChuacus complex and also the Chortıs Block werepart of the Taconic–Acadian orogen along thenorthern margin of South America. The mostobvious correlation, based on lithostratigraphy andtype and age of magmatism, compares the Rabinalcomplex with the Esperanza suite of the Acatlancomplex. In this scenario, at least the pre-Silurianparts of the Rabinal complex should containhigh-P relicts, which are not observed. Placing theRabinal complex (and the Maya Block) furtherinto the hinterland of the Ordovician orogeny thanthe Esperanza suite allows variable metamorphicoverprint. The Chuacus complex best correlateswith the Ixcamilpa suite and the Cosoltepec suiteof the Acatlan complex. These units formed a mar-ginal basin that was transferred into a Silurian–Carboniferous subduction–accretion complex; theChuacus complex may have represented theleading edge of this subduction complex.

After the final amalgamation of Pangaea, an arcdeveloped along its western margin, causing mag-matism and regional amphibolite-facies meta-morphism in southern Mexico (c. 290–250 Ma),the Rabinal complex of the southern Maya Blockand the Chuacus complex (c. 270–215 Ma) andthe Chortıs Block (c. 283–242 Ma). Identical intru-sion and metamorphic ages (273 and 254 Ma) in theChiapas Massif of the southwestern Maya Block andthe northern Chortıs Block suggest a similar pos-ition within the arc. Younger ages in the Rabinaland Chuacus complexes are attributed to their pos-ition in the hinterland of the main arc and a speculat-ive episode of flat-slab subduction.

The Jurassic opening of the Gulf of Mexico sep-arated the Maya Block from North America. Graniteintruded the Rabinal complex (Maya Block) ofwestern Guatemala at 164–177 Ma. Most ofthe Maya Block was probably emergent duringTriassic–Jurassic with the likely Middle Jurassic–Early Cretaceous Todos Santos Formation reflectingrifting during separation of North and SouthAmerica. Dextral shear along c. NNW trendingshear zones within southern Mexico is kinemati-cally related to opening the Gulf of Mexico andmay account for the concentration of magmatismat c. 175–166 Ma in southern Mexico. Granitesalso intruded the northern Chortıs Block duringMiddle Jurassic time (c. 168–159 Ma). TheMiddle Jurassic Agua Frıa Formation recordsupper crustal transtension and rifting. Rifting ofthe Chortıs Block from southern Mexico probably


initiated within the thermally weakened proto-Pacific arc, possibly as a failed arm of the Proto-Caribbean rift. This rifting separated the northernChortıs Block from the Xolapa complex of southernMexico and also parts of the Guerrero terrane fromparts of southwestern Mexico that have Jurassic-aged basement. The Sanarate complex of the north-western Chortıs Block correlates with the Arteaga–Las Ollas units of the Guerrero terrane, sharing adominantly mafic lithology and Jurassic meta-morphism. Rifting produced the Late Jurassic–Early Cretaceous oceanic crust of the South ElTambor complex and terminated before 132 Ma,the age of eclogite-facies metamorphism in thatcomplex.

Rifting between Chortıs Block and southernMexico reverted to subduction in the Early Cretac-eous. Arc magmatism and high-T/low-P meta-morphism (141–126 Ma) along the Xolapacomplex of southern Mexico is probably related tosubduction of South El Tambor complex litho-sphere, which may have been emplaced southwardonto the Chortıs Block at or after c. 120 Ma. Later-ally, Xolapa-complex magmatism is correlated withthat in the Taxco–Viejo unit of the Mixteca terrane.While collision of the Chortıs Block with southernMexico terminated arc magmatism in the Xolapacomplex, proto-Pacific subduction continued onthe Chortıs Block and the Zihuatanejo unit of theGuerrero terrane to its NW. This unit and the Sana-rate complex of the Chortıs Block amalgamated tosouthern Mexico during Late Cretaceous.

Spectacular sinistral transpressive collisionoccurred along the southern margin of the MayaBlock during Late Cretaceous (�76 Ma). Collisiontransferred the southern Maya Block into a nappestack with the North El Tambor complex at top.Sinistral transpression was partitioned on all scalesbetween oblique stretching along the foliation andsinistral strike–slip along steeply dipping shearzones. To first order, deformation is identical inthe North El Tambor complex, Chuacus complex,and the Laramide, northern foreland fold–thrustbelt. The relatively high-P/low-T metamorphismand nearly coeval blueschist- (North El Tamborcomplex) and lower amphibolite-facies metamorph-ism (Chuacus complex) suggest a cold geothermand steep continental subduction. The MayaBlock–North El Tambor complex suture extendswestward into the footwall of the Santa Elena–Nicoya complexes, the Siuna terrane of the southernChortıs Block, and the Nicaragua Rise and to the eastinto the subduction–accretion and arc complexes ofCuba and Hispaniola. Postulated break-off of theNorth El Tambor complex slab may account for�75 Ma magmatism in the southern Maya Block.

High shear stress across the Caribbean–NorthAmerican plate boundary during the Eocene

triggered breakaway of the Chortıs Block fromsouthern Mexico at �40 Ma, coeval with initiationof the Cayman trough pull-apart basin. Sinistralbreak away started in an area, the Xolapa complexand the northern Chortıs Block, that had been anarc at least since Late Cretaceous and thus was ther-mally weakened. The Las Ovejas complex of thenorthern Chortıs Block forms a westward narrowingbelt of distributed sinistral transtension and rep-resents the southern part of the high-strain displace-ment zone along which the Chortıs Block migratedeastward. Deformation interacted with amphibolite-facies metamorphism, widespread migmatization,and local magmatism, illustrated most clearly bysmall syn- to post-tectonic pegmatites. The closespatial and temporal interaction between high-straindeformation and Cenozoic metamorphism, migma-tization and magmatism is confined to the northernChortıs Block. It appears that arc magmatism pro-vided the heat input to localize deformation. Conse-quently, the largest volume and the concentration intime (40–20 Ma) of magmatism and metamorphismseems to have been controlled by shear zonesassociated with transtensional–transpressional tec-tonics at the boundary between southern Mexicoand the Chortıs Block. The Sula graben, and poss-ibly other north-trending grabens of southern Guate-mala and northern Honduras, started in temporal andkinematic compatibility with wrenching along theLas Ovejas complex. The topographically wellexpressed active grabens, however, are neotectonicfeatures. Also the Jocotan–Chamelecon faultzone was part of this deformation system. Break-away and early eastward translation of the ChortısBlock induced exhumation rates of 0.8–1.9 mm/annum in the Xolapa complex and the northernChortıs Block. Initial eastward translation is alsorecorded by rapid cooling within the southern/southwestern Maya Block at c. 30 Ma. Approxi-mately 175, c. 130 and c. 40 Ma magmatic beltsboth in southern Mexico and on the Chortıs Blockconsistently argue for �1100 km offset since�40 Ma along the northern Caribbean Plate bound-ary. Post-40 Ma structures and a c. 15 Ma magmaticbelt indicate that c. 700 km of that offset occurredprior to c. 15 Ma. The Tonala shear zone, boundingthe Chiapas Massif to the SW, and Polochic shearzones were major deformation zones at 9–8 and12–5 Ma, respectively.

North–south rifts in western Honduras andsoutheastern Guatemala are related to plateauuplift following slab break-off underneath CentralAmerica. Apatite–fission track ages of 11.2–7.9 Ma from the western flank of the Sula grabensupport the Late Miocene onset of rifting. Presently,the triple junction between the Cocos, NorthAmerican and Caribbean plates is a zone ofdiffuse deformation. Normal faulting, exemplified


by the grabens in southern Guatemala and westernHonduras, extends northward across the Motaguafault zone into the Chuacus complex, indicatingthat today the Polochic fault may constitute themajor plate boundary fault. Distributed, dextral,WNW trending fault zones are active along theactive volcanic arc.

This work grew out of our research in southern Mexico inthe early 1990s and was started by G. Dengo (GuatemalaCity) in 1993. Stimulating discussions with C. Ranginand his group at College de France, Aix-en-Provence,and the inspiring work of our Mexican colleaguesprompted the first author to put together the data thatalways seemed to be less bright than those from Asia; wehope that this paper proves otherwise. B. Wauschkuhn(Freiberg) compiled Fig. 1b. G. Seward (UCSB) con-tributed the EBSD LPO measurement of sample5PB-10. E. Enkelmann (Lehigh) supported the AFTwork. W. Frisch (Tubingen), G. Michel (Bermuda) andU. Riller (Toronto) helped in the field. L. Ratschbacherthanks B. Hacker (UCSB) for endless support in theAr-lab when both were at Stanford in the mid1990s. M. Kozuch collected sample SF1D, andA. Andrews-Rodbell and J. M. Gutierrez sample86TA100. Funded by DFG grants RA442/6 and 25 (toL. Ratschbacher), NSF grants EAR 0510325-02 (toU. Martens), EAR 9219284 (to M. Gordon), CONACYTproject D41083-F (to B. Weber), and awards fromDAAD and Studienstiftung des Deutschen Volkes toK. Stubner. Reviewed by L. Kennan, J. L. Pindell and ananonymous colleague. Finally, we thank the editors toallow a long and data-packed paper into their volume.

References

ALANIZ-ALVAREZ, S., VAN DER HEYDEN, P., NIETO-SAMANIEGO, A. F. & ORTEGA-GUTIERREZ, F.1996. Radiometric and kinematic evidence forMiddle Jurassic strike–slip faulting in southernMexico related to the opening of the Gulf of Mexico.Geology, 24, 443–446.

ALVAREZ-GOMEZ, J. A., MEIJER, P. T., MARTINEZ-DIAZ, J. J. & CAPOTE, R. 2008. Constraints fromfinite element modeling on the active tectonics of north-ern Central America and the Middle America Trench.Tectonics, 27, TC1008; doi: 10.1029/2007TC002162.

ANDERSON, J. L. & SMITH, D. R. 1995. The effects oftemperature and f O2 on the Al-in-hornblende barom-eter. American Mineralogist, 80, 549–559.

ANDERSON, T. H., BURKART, B., CLEMONS, R. E.,BOHNENBERGER, O. H. & BLOUNT, D. N. 1973.Geology of the western Altos Cuchumatanes, north-western Guatemala. Geological Society of AmericaBulletin, 84, 805–826.

ANDREANI, L., LE PICHON, X., RANGIN, C. &MARTINEZ-REYES, J. 2008. The southern MexicoBlock, main boundaries and new estimation for itsQuaternary motion. Bulletin de la Societe Geologiquede France, 179, 209–223.

AVE LALLEMANT, H. G. & GORDON, M. B. 1999. Defor-mation history of Roatan Island, Implications for the

origin of the Tela Basin (Honduras). In: MANN, P.(ed.) Caribbean Basins. Sedimentary Basins of theWorld, 4. Elsevier, Amsterdam, 197–218.

BANNO, S. 1959. Aegirinaugites from crystalline schistsin Sikoku. Journal Geological Society Japan, 65,652–657.

BARGAR, K. E. 1991. Fluid inclusions and preliminarystudies of hydrothermal alteration in core holePLTG-1, Platanares geothermal area, Honduras.Journal of Volcanology and Geothermal Research,45, 147–160.

BATESON, J. H. 1972. New interpretation of geology ofMaya Mountains, British Honduras. American Associ-ation of Petroleum Geologists, Bulletin, 56, 956–963.

BECCALUVA, L., BELLIA, S. ET AL. 1995. The northwes-tern border of the Caribbean Plate in Guatemala: newgeological and petrological data on the Motagua ophio-litic belt. Ofioliti, 20, 1–15.

BERTRAND, J. & VUAGNAT, M. 1975. Donnees chimi-ques diverses sur des ophiolites du Guatemala. BulletinSuisse de Mineralogie Petrographie, 57, 466–483.

BERTRAND, J., DELALOYE, M., FONTIGNIE, D. &VUAGNAT, M. 1978. Ages (K–Ar) sur diverses ophio-lites et roches associees de la Cordillere centrale duGuatemala. Bulletin Suisse de Mineralogie Petrogra-phie, 58, 405–412.

BGR, BUNDESANSTALT FUR BODENFORSCHUNG 1971.Geologische Ubersichtskarte 1:125 000, BajaVerapaz und Sudteil der Alta Verapaz (Guatemala).D-3 Hannover-Buchholz.

BLUNDY, J. D. & HOLLAND, T. 1990. Calcic amphiboleequilibria and a new plagioclase-amphibole geo-thermometer. Contributions to Mineralogy and Petro-logy, 104, 204–224.

BOHLEN, S. & LIOTTA, J. 1986. A barometer for garnetamphibolites and garnet granulites. Journal of Petrol-ogy, 27, 1025–1034.

BONIS, S., BOHNENBERGER, O. H. & DENGO, G. 1970.Mapa geologio de la Republica de Guatemala. Insti-tuto Geografico Nacional, Guatemala.

BRUECKNER, H. K., HEMMING, S., SORENSEN, S. &HARLOW, G. E. 2005. Synchronous Sm-Nd mineralages from HP terranes on both sides of the MotaguaFault of Guatemala: convergent suture and strike–slip fault? AGU Fall Meeting, Supplement, 86, 52.

BURKART, B. 1983. Neogene North America–CaribbeanPlate boundary across northern Central America:offset along the Polochic fault. Tectonophysics, 99,251–270.

BURKART, B. 1994. Northern Central America. In:DONOVAN, S. & JACKSON, T. (eds) CaribbeanGeology. UWI Publ. Association, Kingston, 265–283.

BURKART, B., DEATON, B. C., DENGO, C. & MORENO,G. 1987. Tectonic wedges and offset Laramide struc-tures along the Polochic fault of Guatemala andChiapas, Mexico: reaffirmation and large Neogene dis-placement. Tectonics, 6, 411–422.

CAMPA, M. F. & IRIONDO, A. 2004. Significado de data-ciones cretacicas de los arcos volcanicos de Taxco,Taxco Viejo y Chacalapa, en la evolucion de la plata-forma Guerrero Morelos. IV Reunion Nacional deCiencias de la Tierra, Juriquilla, Queretaro, 338.

CARSWELL, D., O’BRIAN, P., WILSON, R. & ZHAI, M.2000. Metamorphic evolution, mineral chemistry and


thermobarometry of schists and orthogneisses hostingultra-high pressure eclogites in the Dabie Shan ofCentral China. Lithos, 52, 121–155.

CENTENO-GARCIA, E. 2005. Review of Upper Paleozoicand Lower Mesozoic stratigraphy and depositionalenvironments of central and west Mexico: constraintson terrane analysis and paleogeography. In: ANDER-

SON, T. H., NOURSE, J. A., MCKEE, J. W. &STEINER, M. B. (eds) The Mojave-Sonora MegashearHypothesis: Development, Assessment, and Alterna-tives. Geological Society of America Special Paper,393, 233–258.

CENTENO-GARCIA, E., RUIZ, J., CONEY, P. J.,PATCHETT, P. J. & ORTEGA, G. F. 1993. Guerreroterrane of Mexico: its role in the Southern Cordillerafrom new geochemical data. Geology, 21, 419–422.

CERCA, M., FERRARI, L., LOPEZ-MARTINEZ, M.,MARTINY, B. & IRONDO, A. 2007. Late Cretaceousshortening and Early Cenozoic shearing in the centralSierra Madre des Sur, southern Mexico: insights intothe evolution of the Caribbean–North Americanplate interaction. Tectonics, 26, TC3007; doi:10.1029/2006TC001981.

CHIARI, M., DUMITRICA, P., MARRONI, M., PANDOLFI,L. & PRINCIPI, G. 2006. Radiolarian biostratigraphicevidence for a Late Jurassic age of the El TamborGroup ophiolites (Guatemala). Ofioliti, 31, 173–182.

CLEMONS, R. E. 1966. Geology of the Chiquimulaquadrangle, Guatemala, Central America. PhDthesis, University of Texas, Austin, TX.

COLOMBI, A. 1988. Metamorphisme et geochimie desroches mafigues des Alpes onest centrals (geoprofilViege–Domodossola–Lacarno. PhD thesis, Univer-sity Lausanne, Switzerland.

CORONA-CHAVEZ, P., POLI, S. & BIGIOGGERO, B. 2006.Syn-deformational migmatites and magmatic-arcmetamorphism in the Xolapa Complex, southernMexico. Journal of Metamorphic Geology, 24,169–191.

DAHL, P. S. & FRIBERG, L. M. 1980. The occurrence andchemistry of epidote-clinozoisites in mafic gneissesfrom the Ruby Range, southwestern Montana. Contri-butions to Geology, University Wyoming, 18, 77–82.

DEMETS, C. 2001. A new estimate for present-dayCocos–Caribbean Plate motion: implications for slipalong the Central American volcanic arc. GeophysicalResearch Letters, 28, 4043–4046.

DENGO, G. 1969. Problems of tectonic relations betweenCentral America and the Caribbean. Transactions ofthe Gulf Coast Association of Geological Societies,19, 311–320.

DICKINSON, W. R. & LAWTON, T. F. 2001. Carboniferousto Cretaceous assembly and fragmentation of Mexico.Geological Society of America Bulletin, 113,1142–1160.

DONELICK, R. A., RODEN, M. K., MOOERS, J. D.,CARPENTER, B. S. & MILLER, D. S. 1990. Etchablelength reduction of induced fission tracks in apatite atroom temperature (¼23 8C): crystallographic orien-tation effects and ‘initial’ mean lengths. NuclearTracks and Radiation Measurements, 17, 261–265.

DONNELLY, T. W., HORNE, G. S., FINCH, R. C. &LOPEZ-RAMOS, E. 1990. Northern Central America:the Maya and Chortis blocks. In: DENGO, G. &

CASE, J. (eds) The Caribbean Region. Geology ofNorth America. Geological Society of America, H,37–76.

DROBE, J. & OLIVER, D. 1998. U–Pb age constraints onEarly Cretaceous volcanism and stratigraphy incentral Honduras, Geological Society of America.GSA Cordilleran Section Meeting, 30, 12.

DUCEA, M., GEHRELS, G. E., SHOEMAKER, S., RUIZ, J.& VALENCIA, V. A. 2004a. Geologic evolution ofthe Xolapa Complex, southern Mexico: evidencefrom U–Pb zircon geochronology. GeologicalSociety of America Bulletin, 116, 1016–1025.

DUCEA, M., VALENCIA, V. A. ET AL. 2004b. Rates ofsediment recycling beneath the Acapulco trench:constraints from (U-Th)/He thermochronology.Journal of Geophysical Research, 109; doi: 10.1029/2004JB003112.

ELIAS-HERRERA, M. & ORTEGA-GUTIERREZ, F. 2002.Caltepec fault zone: an Early Permian dextral transpres-sional boundary between the Proterozoic Oaxacan andPalaeozoic Acatlan complexes, southern Mexico andregional tectonic implications. Tectonics, 21, 1–19.

ELIAS-HERRERA, M., ORTEGA-GUTIERREZ, F.,SANCHEZ-ZAVALA, J. L., MACIAS-ROMO, C.,ORTEGA-RIVERA, A. & IRIONDO, A. 2005. La fallade Caltepec: raıces expuestas de una frontera tectonicade larga vida entre dos terrenos continentales del sur deMexico. Boletın de la Sociedad Geologica Mexicana,LVII, 83–109.

EMMET, P. 1983. Geology of the Agalteca Quadrangle,Honduras, Central America. MS thesis, The Universityof Texas, Austin.

ESCUDER VIRUETE, J., CONTRERAS, F. ET AL. 2006a.Transpression and strain partitioning in the Caribbeanisland-arc: fabric development, kinematics and Ar-Arages of syntectonic emplacement of the Loma deCabrera batholith, Dominican Republic. Journal ofStructural Geology, 28, 1496–1519.

ESCUDER VIRUETE, J., DIAZ DE NEIRA, A. ET AL. 2006b.Magmatic relationships and ages of Caribbean Islandarc tholeiites, boninites and related felsic rocks,Dominican Republic. Lithos, 90, 161–186.

FERRARI, L., GARDUNO, V. H., PASQUARE, G. &TIBALDI, A. 1994. Volcanic and tectonic evolutionof central Mexico: oligocene to present. GeofısicaInternacional, 33, 91–105.

FLORES, K., BAUMGARTNER, P. O., SKORA, S.,BAUMGARTNER, L., COSCA, M. & CRUZ, D. 2007.The Siuna serpentinite melange: an Early Cretaceoussubduction/accretion of a Jurassic arc. EOS Trans-actions, AGU Fall Meeting, Supplement, 88/52,Abstract T11D-03.

FOURCADE, E., MENDEZ, J. ET AL. 1994. Dating of thesettling and drowning of the carbonate platform, andof the overthrusting of the ophiolites on the MayaBlock during the Mesozoic (Guatemala). NewsletterStratigraphy, 30, 33–43.

FOURCADE, E., PICCIONI, L., ESCRIBA, J. & ROSSELO,E. 1999. Cretaceous stratigraphy and palaeoenviron-ments of the Southern Peten Basin, Guatemala. Cretac-eous Research, 20, 793–811.

FUHRMAN, M. L. & LINDSLEY, D. H. 1988. Ternaryfeldspar modeling and thermometry. American Miner-alogist, 73, 201–215.


GANGULY, J., CHENG, W. & TIRONE, M. 1996. Thermo-dynamics of aluminosilicate garnet solid solution: newexperimental data, an optimized model, and thermo-metry applications. Contributions to Mineralogy andPetrology, 126, 137–151.

GARCIA-CASCO, A., TORRES-ROLDAN, R. L., MILLAN,G., MONIE, P. & SCHNEIDER, J. 2002. Oscillatoryzoning of eclogitic garnet and amphibole, NorthernSerpentinite melange, Cuba: a record of tectonicinstability during subduction. Journal of MetamorphicGeology, 20, 581–598.

GARCIA-CASCO, A., LAZARO, C., ROJAS AGRAMONTE,Y., KRONER, A. & NEUBAUER, F. 2007. From Aptianonset to Danian demise of subduction along the north-ern margin of the Caribbean Plate (Sierra del ConventoMelange, Eastern Cuba). AGU Fall Meeting, Sup-plement, 88/52, Abstract T11D-06.

GHENT, E. D., STOUT, M. Z., BLACK, P. M. & BROTH-

ERS, R. N. 1987. Chloritoid bearing rocks with blues-chists and eclogites, northern New Caledonia. Journalof Metamorphic Petrology, 5, 239–254.

GOMBERG, D. N., BANKS, P. O. & MCBIRNEY, A. R.1968. Guatemala: preliminary zircon ages from theCentral Cordillera. Science, 162, 121–122.

GORDON, M. B. 1993. Revised Jurassic and Early Cretac-eous (pre-Yojoa Group) stratigraphy of the Chortisblock: paleogeographic and tectonic implications.GCSSEPM Foundation 13th Annual Research Confer-ence, Proceedings, 143–154.

GORDON, M. B. & MUEHLBERGER, W. R. 1994. Rotationof the Chortis block causes dextral slip of the Guayapefault. Tectonics, 13, 858–872.

GOSE, W. A. 1985. Paleomagnetic results from Hondurasand their bearing on Caribbean tectonics. Tectonics, 4,565–585.

GRAFE, F., STANEK, K. P. ET AL. 2001. Rb–Sr and40Ar–39Ar mineral ages of granitoid intrusives in theMabujina unit, Central Cuba: thermal exhumationhistory of the Escambray Massif. Journal of Geology,109, 615–631.

GRAHAM, C. M. & POWELL, R. 1986. A garnet-hornblende geothermometer: calibration, testing, andapplication to the Pelona Schists, Southern California.Journal of Metamorphic Geology, 2, 13–31.

GRAY, G. G., POTTORF, R. J., YUREWICZ, D. A.,MAHON, K. I., PEVEAR, D. R. & CHUCHLA, R. J.2001. Thermal and chronological record of syn- topost-Laramide burial and exhumation, Sierra MadreOriental, Mexico. In: BARTOLINI, C., BUFFLER,R. T. & CANTU-CHAPA, A. (eds) The Western Gulfof Mexico Basins: Tectonics, Sedimentary Basins,and Petroleum Systems. American Association ofPetroleum Geologists, Memoirs, 75, 159–181.

GREEN, T. H. & HELLMAN, P. L. 1982. Fe–Mg partition-ing between coexisting garnet and phengite at highpressure, and comments on a garnet–phengitegeothermometer. Lithos, 15, 253–266.

GUZMAN-SPEZIALE, M. 2001. Active seismic defor-mation in the grabens of northern Central Americaand its relationship to the relative motion of theNorth America–Caribbean Plate boundary. Tectono-physics, 337, 39–51.

HALL, C. M., KESLER, S. E. ET AL. 2004. Age and tectonicsetting of the Camaguey volcanic-intrusive arc,

Cuba: evidence for rapid uplift of the western GreaterAntilles. Journal of Geology, 112, 521–542.

HARLOW, G. E. 1994. Jadeitites, albitites and related rocksfrom the Motagua Fault Zone, Guatemala. Journal ofMetamorphic Geology, 12, 49–68.

HARLOW, G. E., HEMMING, S. R., AVE LALLEMANT,H. G., SISSON, V. B. & SORENSEN, S. S. 2004. Twohigh-pressure–low-temperature serpentinite-matrixmelange belts, Motagua fault zone, Guatemala: arecord of Aptian and Maastrichtian collisions.Geology, 32, 17–20.

HERNANDEZ-TREVINO, T., TORRES DE LEON, R.,SOLIS-PICHARDO, G., SCHAAF, P., HERNANDEZ-BERNAL, M. S. & MORALES-CONTRERAS, J. J.2004. Edad de la Formacion Chapolapa en la localidaddel Rıo Cochoapa al este del Ocotito, estado deGuerrero. IV Reunion Mexicana de Ciencias de laTierra, Juriquilla, Queretaro, 338.

HERRMANN, U., NELSON, B. K. & RATSCHBACHER, L.1994. The origin of a terrane: U–Pb systematics andtectonics of the Xolapa complex (southern Mexico).Tectonics, 13, 455–474.

HEY, M. 1954. A new review of the chlorites, Mineralogi-cal Magazine, 30, 225–234.

HODGES, K. V. & SPEAR, F. S. 1982. Geothermometry,geobarometry and the Al2SiO5 triple point atMt. Moosilauke, New Hampshire. American Mineral-ogist, 67, 1118–1134.

HOERNLE, K. & HAUFF, F. 2005. Oceanic igneous com-plexes. In: BUNDSCHUH, J. & ALVARADO, G. (eds)Central America: Geology, Resources, and NaturalHazards. Taylor and Francis, Abingdom, 523–547.

HOLLAND, T. J. B. 1980. The reaction albite ¼ jadeiteþquartz determined experimentally in the range of600 8C–1200 8C. American Mineralogist, 65,129–134.

HOLLAND, T. J. B. 1990. Activities of components inomphacitic solid solution. Contributions MineralogyPetrology, 105, 446–453.

HOLLAND, T. J. B. & BLUNDY, J. D. 1994. Non-idealinteractions in calcic amphiboles and their bearing onamphibole-plagioclase thermometry. Contributions toMineralogy and Petrology, 116, 208–224.

HOLLAND, T. J. B. & POWELL, R. 1998. An internallyconsistent thermodynamic data set for phases of petro-logical interest. Journal of Metamorphic Geology, 16,309–343.

HORNE, G., ATWOOD, M. & KING, A. 1974. Stratigraphy,sedimentology, and paleoenvironment of EsquiasFormation of Honduras. American Association ofPetroleum Geologists Bulletin, 58, 176–188.

HORNE, G., CLARK, G. & PUSHKAR, P. 1976a. Pre-Cretaceous rocks of northwestern Honduras: basementterrane in Sierra de Omoa. American Association ofPetroleum Geologists Bulletin, 60, 566–583.

HORNE, G., PUSHKAR, A. & SHAFIQULLAH, M. 1976b.Laramide plutons on the landward continuation of theBonacca ridge, northern Honduras. PublicacionesGeologicas del ICAITI (Guatemala), 5, 84–90.

IGN. 1978. Instituto Geografico Nacional, Mapa Geolo-gico de Guatemala. Escala 1 : 50 000, Hoja RıoHondo, Guatemala.

INDARES, A. & MARTIGNOLE, J. 1985. Biotite garnetgeothermometry in the granulite facies: the influence


of Ti and Al in biotite. American Mineralogist, 70,272–278.

JAMES, K. H. 2006. Arguments for and against the Pacificorigin of the Caribbean Plate: discussion, findingfor an inter-American origin. Geological Acta, 4,279–302.

JONCKHEERE, R. 2003a. On the densities of etchablefission tracks in a mineral and co-irradiated externaldetector with reference to fission-track dating of min-erals. Chemical Geology, 200, 41–58.

JONCKHEERE, R. 2003b. On the methodical problems inestimating geological temperature and time frommeasurements of fission tracks in apatite. RadiationMeasurements, 36, 43–55.

JONCKHEERE, R. & WAGNER, G. A. 2000. The KTBapatite fission-track profile: the significance of bore-hole data in fission-track analysis. Geological Societyof Australia Abstracts, 58, 193–194.

JORDAN, B. R., SIGURDSSON, H. ET AL. 2007.Petrogenesisof Central American Cenozoic ignimbrites andassociated Caribbean Sea tephra. In: MANN, P. (ed.)Geologic and Tectonic Development of the CaribbeanPlate Boundary in Northern Central America. Geo-logical Society of America, Special Papers, 428,151–179.

KEPPIE, D. J. 2004. Terranes of Mexico revisited: a 1.3billion year odyssey. International Geology Review,46, 765–794.

KEPPIE, D. J., NANCE, R. D. ET AL. 2004. Mid-Jurassictectonothermal event superposed on a Palaeozoic geo-logical record in the Acatlan complex of southernMexico: hotspot activity during the breakup ofPangea. Gondwana Research, 7, 239–260.

KEPPIE, J. D., NANCE, R. D., FERNANDEZ-SUAREZ, J.,STOREY, C. D., JEFFRIES, T. E. & MURPHY, J. B.2006. Detrital zircon data from the eastern Mixtecaterrane, southern Mexico: evidence for anOrdovician–Mississippian continental rise and aPermo-Triassic clastic wedge adjacent to Oaxaquia.International Geology Review, 48, 97–111.

KESLER, S. E., CAMPELL, I. H. & ALLEN, CH. M. 2005.Age of the Los Ranchos Formation, Dominican Repub-lic: timing and tectonic setting of primitive island arcvolcanism in the Caribbean region. GeologicalSociety of America Bulletin, 117, 987–995.

KETCHAM, R. A., DONELICK, R. A. & DONELICK, M. B.2000. AFTSolve: a program for multikinetic modelingof apatite fission-track data. Geological MaterialsResearch, 2, 1–32.

KOHN, M. & SPEAR, F. 1990. Two new geobarometers forgarnet-amphibolites, with application to southeasternVermont. American Mineralogist, 75, 89–96.

KREBS, M., STANEK, K. P. ET AL. 2007. Age of highpressure metamorphism from the Escambray Massif,Cuba. Abstract, Goldschmidt Conference, Cologne,2007. Cambridge Publications, A522.

KRETZ, R. 1983. Symbols for rock-forming minerals.American Mineralogist, 68, 277–279.

KROGH, E. J. 2000. The garnet-clinopyroxene Fe/Mggeothermometer: an updated calibration. Journal ofMetamorphic Geology, 18, 211–219.

KROGH, E. J. & TERRY, M. P. 2004. Geothermobarometryof UHP and HP eclogites and schists: an evaluation

of equilibria among garnet-clinopyroxene-kyanite-phengite-coesite/quartz. Journal of MetamorphicGeology, 22, 579–592.

LEAKE, B. E., WOOLEY, A. R. ET AL. 1997. Nomenclatureof amphiboles. Report of the subcommittee on amphi-boles of the Intern. Mineral. Association Commissionon new minerals and mineral names. EuropeanJournal of Mineralogy, 9, 623–665.

LUDWIG, K. R. 2003. Isoplot 3.01. A GeochronologicalToolkit for Microsoft Exel. Berkeley GeochronologyCenter Special Publications, 4, 1–70.

MACHORRO, R. 1993. Geology of the northwestern partof the Granados Quadrangle, Central Guatemala.MSc thesis, University of Texas, El Paso.

MALONE, J. R., NANCE, R. D., KEPPIE, J. D. & DOSTAL,J. 2002. Deformational history of part of the AcatlanComplex: Late Ordovician Early Permian orogenesisin southern Mexico. Journal of South AmericanEarth Sciences, 15, 511–524.

MANN, P. 2007. Overview of the tectonic history ofnorthern Central America. In: MANN, P. (ed.) Geo-logic and Tectonic Development of the CaribbeanPlate Boundary in Northern Central America.Geological Society of America, Special Papers, 428,1–19.

MANTON, W. I. 1996. The Grenville of Honduras. Geo-logical Society of America, Annual Meeting Abstractswith Program, A-493.

MANTON, W. I. & MANTON, R. S. 1984. Geochronologyand Late Cretaceous–Tertiary Tectonism of Hon-duras. Direccion General de Minas e Hidrocarburos,Honduras, 1–55.

MANTON, W. I. & MANTON, R. S. 1999. The southernflank of the Tela basin, Republic of Honduras. In:MANN, P. (ed.) Caribbean Basins. SedimentaryBasins of the World, 4. Elsevier, Amsterdam, 219–236.

MARTENS, U., ORTEGA-OBREGON, C., ESTRADA, J. &VALLE, M. 2007. Metamorphism and metamorphicrocks. In: BUNDSCHUH, J. & ALVARADO, G. (eds)Central America: Geology, Resources, and NaturalHazards. Taylor and Francis, Abingdom, 485–522.

MARTENS, U., WEBER, B. & VALENCIA, V. 2010. U–Pbgeochronology of lower Paleozoic beds in the south-western Maya block, Central America: its affinitywith Avalonian-type Peri-Gondwanan terranes. Geo-logical Society of America Bulletin, in press.

MASSONNE, H.-J. & SZPURKA, Z. 1997. Thermodynamicproperties of white micas on the basis of high-pressureexperiments in the system K2O–MgO–Al2O3–SiO2–H2O and K2O–FeO–Al2O3–SiO2–H2O. Lithos, 41,229–250.

MESCHEDE, M. & FRISCH, W. 1998. A plate-tectonicmodel for the Mesozoic and Early Cenozoichistory of the Caribbean Plate. Tectonophysics, 296,269–291.

MESCHEDE, M., FRISCH, W., HERRMANN, U. &RATSCHBACHER, L. 1996. Stress transmission acrossan active plate boundary: an example from southernMexico. Tectonophysics, 266, 81–100.

MMAJ, METAL MINING AGENCY OF JAPAN. 1980.Report on Geology Survey of the Western Area,Olancho. Japan International Cooperation Agency,Government of Japan, 1–138.


MORAN-ZENTENO, D. 1992. Investigaciones isotopicasde Rb–Sr y Sm–Nd en rocas cristalinas de laeregion de Tierra Colorada–Acapulco–Cruz–Grande, Estado de Guerrero. PhD thesis, UniversidadAutonoma de Mexico, Mexico, DF.

MORAN-ZENTENO, D. J., CORONA-CHAVEZ, P. &TOLSON, G. 1996. Uplift and subduction erosion insouthwestern Mexico since the Oligocene: Pluton geo-barometry constraints. Earth and Planetary ScienceLetters, 141, 51–65.

MORAN-ZENTENO, D. J., TOLSON, G. ET AL. 1999.Tertiary arc-magmatism of the Sierra Madre del Sur,Mexico, and its transition to the volcanic activity ofthe Trans-Mexican Volcanic Belt. Journal of SouthAmerican Earth Sciences, 12, 513–535.

NIETO-SAMANIEGO, A. F., ALANIZ-ALVAREZ, S. A.,SILVA-ROMO, G., EGUIZA-CASTRO, M. H. &MENDOZA-ROSALES, C. C. 2006. Latest Cretaceousto Miocene deformation events in the eastern SierraMadre del Sur, Mexico, inferred from the geometryand age of major structures. Geological Society ofAmerica Bulletin, 118, 238–252.

ORTEGA-GUTIERREZ, F., ELIAS-HERRERA, M.,MACIAS-ROMERO, C. & LOPEZ, R. 1999. Late Ordo-vician–Early Silurian continental collisional orogenyin southern Mexico and its bearing on Gondwana–Laurentia connections. Geology, 27, 719–722.

ORTEGA-GUTIERREZ, F., SOLARI, L. A., SOLE, J.,MARTENS, U., GOMEZ-TUENA, A., MORAN-ICAL,S., REYES-SALA, M. & ORTEGA-OBREGON, C.2004. Polyphase, high-temperature eclogite-faciesmetamorphism in the Chuacus complex, CentralGuatemala: petrology, geochronology, and tectonicimplications. International Geology Review, 46,445–470.

ORTEGA-GUTIERREZ, F., SOLARI, L. A. ET AL. 2007. TheMaya–Chortıs boundary: a tectonostratigraphicapproach. International Geology Review, 49,996–1024.

ORTEGA-OBREGON, C., SOLARI, L., ORTEGA-GUTIERREZ, F., SOLE-VINAS, J. & GOMEZ-TUENA,A. 2004. Caracterizacion estructural, petrologica ygeocronologica de la zona de cizalla “Baja Verapaz”,Guatemala. Libro de Resumenes, IV Reunion Nacionalde Ciencias de la Tierra, Juriquilla, 2004, 204.

ORTEGA-OBREGON, C., SOLARI, L., KEPPIE, J. D.,ORTEGA-GUTIERREZ, F., SOLE, J. & MORAN, S.2008. Middle–Late Ordovician magmatism andLate Cretaceous collision in the southern Mayablock, Rabinal-Salama area, central Guatemala:implications for North America–Caribbean Platetectonics. Geological Society of America Bulletin,120, 556–570.

PERCHUK, L. L. & LAVRENT’EVA, J. V. 1983. Exper-imental investigations of exchange equilibria in thesystem cordierite–garnet–biotite. In: SAXENA, S.(ed.) Kinematics and Equilibria in Mineral Reactions.Springer, Berlin, 199–239.

PINDELL, J. K., CANDE, S. C. ET AL. 1988. A plate-kinematic framework for models of Caribbean evol-ution. Tectonophysics, 155, 121–138.

PINDELL, J. L., KENNAN, L., STANEK, K. P., MARESCH,W. & DRAPER, G. 2006. Foundations of Gulf of

Mexico and Caribbean evolution: eight controversiesresolved. Geologica Acta, 99, 303–341.

PLAFKER, G. 1976. Tectonic aspects of the Guatemalaearthquake of 4 February 1976. Science, 193,1201–1208.

PLYUSNINA, L. P. 1982. Geothermometry and geobaro-metry of plagioclase-hornblende bearing assemblages.Contributions Mineralogy Petrology, 80, 140–146.

RATSCHBACHER, L., RILLER, U., MESCHEDE, M.,HERRMANN, U. & FRISCH, W. 1991. Second look atsuspect terranes in southern Mexico. Geology, 19,1233–1236.

RILLER, U., RATSCHBACHER, L. & FRISCH, W. 1992.Left-lateral transtension along the Tierra Coloradadeformation zone, northern margin of the Xolapamagmatic arc of southern Mexico. Journal of SouthAmerican Earth Sciences, 5, 237–249.

ROBINSON, K. L., GASTIL, R. G., CAMPA, M. F. &RAMIREZ-ESPINOSA, J. 1989. Geochronology ofbasement and metasedimentary rocks in southernMexico and their relation to metasedimentary rocksin Peninsular California. Geological Society ofAmerica Abstracts with Programs, 21, 135.

ROGERS, R. D. 2003. Jurassic–Recent tectonic and strati-graphic history of the Chortis block of Honduras andNicaragua (northern Central America). PhD thesis,University of Texas, Austin, TX.

ROGERS, R. D., KARASON, H. & VAN DER HILST, R. D.2002. Epeirogenic uplift above a detached slab innorthern Central America. Geology, 30, 1031–1034.

ROGERS, R. D., MANN, P., SCOTT, R. W. & PATINO, L.2007a. Cretaceous intra-arc rifting, sedimentation,and basin inversion in east-central Honduras. In:MANN, P. (ed.) Geologic and Tectonic Developmentof the Caribbean Plate Boundary in NorthernCentral America. Geological Society of America,Special Papers, 428, 89–128.

ROGERS, R. D., MANN, P., EMMET, P. A. & VENABLE,M. E. 2007b. Colon fold belt of Honduras: evidencefor Late Cretaceous collision between the continentalChortis block and intra-oceanic Caribbean Arc. In:MANN, P. (ed.) Geologic and Tectonic Developmentof the Caribbean Plate Boundary in NorthernCentral America. Geological Society of America,Special Papers, 428, 129–149.

ROGERS, R. D., MANN, P. & EMMET, P. A. 2007c. Tec-tonic terranes of the Chortis block based on integrationof regional aeromagnetic and geologic data. In: MANN,P. (ed.) Geologic and Tectonic Development of theCaribbean Plate Boundary in Northern CentralAmerica. Geological Society of America, SpecialPapers, 428, 65–88.

ROJAS-AGRAMONTE, Y., KRONER, A., GARCIA-CASCO, A., ITURRALDE-VINENT, M. A., WINGATE,M. T. D. & LIU, D. Y. 2006. Geodynamic implicationsof zircon ages from Cuba. Geophysical ResearchAbstracts, 8, 04943; SRef-ID: 1607-7962/gra/EGU06-A-04943.

ROSENCRANTZ, E., ROSS, M. I. & SLATER, J. G. 1988.Age and spreading history of the Cayman Trough asdetermined from depth, heat flow, and magneticanomalies. Journal Geophysical Research, 93,2141–2157.


RYBURN, R. J., RAHEIM, A. & GREEN, D. H. 1976. Deter-mination of the PT paths of natural eclogites duringmetamorphism – record of a subduction. Lithos, 9,161–164.

SANCHEZ-ZAVALA, J. L., ORTEGA-GUTIERREZ, F.,KEPPIE, J. D., JENNER, G. A. & BELOUSOVA, E.2004. Ordovician and Mesoproterozoic zircon fromthe Tecomate Formation and Esperanza granitoid,Acatlan complex, southern Mexico: local provenancein the Acatlan and Oaxacan complexes. InternationalGeology Review, 46, 1005–1021.

SCHAAF, P., MORAN-ZENTENO, D., HERNANDEZ-BERNAL, M. S., SOLIS-PICHARDO, G., TOLSON, G.& KOHLER, H. 1995. Paleogene continental margintruncation in southwestern Mexico: geochronologicalevidence. Tectonics, 14, 1339–1350.

SCHAAF, P., WEBER, B., WEIS, P., GROß, A., ORTEGA-GUTIERREZ, F. & KOHLER, H. 2002. The ChiapasMassif (Mexico) revised: new geologic and isotopicdata for basement characteristics. Neues JahrbuchGeologie Palaontologie, Abhandlungen, 225, 1–23.

SCHMID, R., FRANZ, L., OBERHANSLI, R. & DONG, S.2000. High Si-phengite, mineral chemistry and P-Tevolution of ultra-high-pressure eclogites and calc-silicates from the Dabie Shan, eastern PR China.Geological Journal, 35, 185–207.

SCHNEIDER, J., BOSCH, D. ET AL. 2004. Origin and evol-ution of the Escambray Massif (Central Cuba): anexample of HP/LT rocks exhumed during intraoceanicsubduction. Journal of Metamorphic Geology, 22,227–247.

SCHWARTZ, D. P. 1976. Geology of the Zacapa quadran-gle and vicinity, Guatemala, Central America. PhDthesis, State University of New York, Binghamton.

SOLARI, L. A., DOSTAL, J., ORTEGA-GUTIERREZ, F. &KEPPIE, J. D. 2001. The 275 Ma arc-related LaCarbonara stock in the northern Oaxacan Complex ofsouthern Mexico: U–Pb geochronology and geochem-istry. Revista Mexicana de Ciencias Geologicas, 18,149–161.

SOLARI, L. A., KEPPIE, J. D., ORTEGA-GUTIEREZ, F.,ORTEGA-RIVERA, A., HAMES, W. E. & LEE,J. K. W. 2004. Phanerozoic structures in the Grenvillecomplex, southern Mexico: result of thick-skinned tec-tonics. International Geology Review, 46, 614–628.

SOLARI, L. A., TORRES DE LEON, R., HERNANDEZ

PINEDA, G., SOLE, J., SOLIS-PICHARDO, G. &HERNANDEZ-TREVINO, G. 2007. Tectonic signifi-cance of Cretaceous–Tertiary magmatic and structuralevolution of the northern margin of the XolapaComplex, Tierra Colorada area, southern Mexico.Geological Society of America Bulletin, 119,1265–1279.

SOLARI, L. A., ORTEGA-GUTIERREZ, F. ET AL. 2009.U–Pb zircon geochronology of Palaeozoic units inWestern and Central Guatemala: insights into the tec-tonic evolution of Middle America. In: JAMES, K. H.,LORENTE, M. A. & PINDELL, J. L. (eds) TheOrigins and Evolution of the Caribbean Plate. Geo-logical Society, London, Special Publications, 328,295–313.

SPEAR, F. S. 1981. NaSi–CaAl exchange equilibriumbetween plagioclase and amphibole. Contributions toMineralogy and Petrology, 72, 33–41.

STANEK, K. P., RIßE, A., RENNO, A., ROMER, R. &GRAFE, F. 2005. The history of the Great AntilleanIsland Arc: example from Central Cuba. Abstract,Colloquium on Latin American Geosciences 2005,Potsdam, Terra Nostra, 1.

STEINER, M. B. & WALKER, J. D. 1996. Late Silurianplutons in Yucatan. Journal Geophysical Research,101, 17727–17735.

STIPP, M., STUNITZ, H., HEILBRONNER, R. & SCHMID,S. M. 2002. The eastern Tonale fault zone: a naturallaboratory for crystal plastic deformation of quartzover a temperature range from 250 to 7008C. Journalof Structural Geology, 24, 1861–1884.

SUTTER, J. F. 1979. Late Cretaceous collisional tectonicsalong the Motagua fault zone, Guatemala. GeologicalSociety of America, Abstracts with Programs, 11,525–526.

TALAVERA-MENDOZA, O., RUIZ, J., GEHRELS, G. E.,MEZA-FIGUEROA, D., VEGA-GRANILLO, R. &CAMPA-URANGA, M. F. 2005. U–Pb geochronologyof the Acatlan complex and implications for thePaleozoic paleogeography and tectonic evolution ofsouthern Mexico. Earth and Planetary ScienceLetters, 235, 682–699.

TALAVERA-MENDOZA, O., RUIZ, J., GEHRELS, G. E.,VALENCIA, V. A. & CENTENO-GARCIA, E. 2007.Detrital zircon U–Pb geochronology of southernGuerrero and western Mixteca arc successions(southern Mexico): new insights for the tectonic evol-ution of southwestern North America during the LateMesozoic. Geological Society of America Bulletin,119, 1052–1065.

TOLSON, G. 2005. La falla Chacalapa en el sur de Oaxaca.Boletın de la Sociedad Geologica Mexicana, 57,111–122.

TORRES, R., RUIZ, J., PATCHETT, P. J. & GRAJALES, J.M. 1999. Permo-Triassic continental arc in easternMexico: tectonic implications for reconstructions ofsouthern North America. In: BARTOLINI, C. W.,WILSON, J. L. & LAWTON, T. F. (eds) Mesozoic Sedi-mentary and Tectonic History of North-CentralMexico. Geological Society of America, SpecialPapers, 340, 191–196.

TRIBOULET, C. 1992. The (Na–Ca)amphibole–albite–chlorite–epidote–quartz geothermobarometer in thesystem S–A–F–M–C–N–H2O. 1. An empiricalcalibration. Journal of Metamorphic Geology, 10,545–556.

TSUJIMORI, T., LIOU, J. G. & COLEMAN, R. G. 2004.Comparison of two contrasting eclogites from theMotagua fault zone, Guatemala: southern lawsoniteeclogite versus northern zoisite eclogite. GeologicalSociety of America Abstracts with Programs, 36, 136.

TSUJIMORI, T., LIOU, J. G. & COLEMAN, R. G. 2005.Coexisting retrograde jadeite and omphacite in ajadeite-bearing lawsonite eclogite from the MotaguaFault Zone, Guatemala. American Mineralogist, 90,836–842.

TSUJIMORI, T., SISSON, V. B., LIOU, J. G., HARLOW,G. E. & SORENSEN, S. S. 2006. Petrologic characteriz-ation of Guatemalan lawsonite eclogite: Eclogitizationof subducted oceanic crust in a cold subduction zone.In: HACKER, B. R., MCCLELLAND, W. C. & LIOU,J. G. (eds) Ultrahigh-pressure Metamorphism: Deep


Continental Subduction. Geological Society ofAmerica, Special Papers, 403, 147–168.

VAN DEN BOOM, G. 1972. Petrofazielle Gliederung desmetamorphen Grundgebirges in der Sierra deChuacus, Guatemala. Beihefte Geologisches Jahrbuch,122, 5–49.

VEGA-GRANILLO, R., TALAVERA-MENDOZA, O. ET AL.2007. Pressure–temperature–time evolution ofPaleozoic high-pressure rocks of the AcatlanComplex (southern Mexico): implications for the evol-ution of the Iapetus and Rheic Oceans. GeologicalSociety of America Bulletin, 119, 1249–1264.

VENABLE, M. 1994. A geological, tectonic, and metallo-genetic evaluation of the Siuna terrane (Nicaragua).PhD thesis, University of Arizona, Tucson.

VILAND, J., HENRY, B., CALIX, R. & DIAZ, C. 1996. LateJurassic deformation in Honduras: proposals for arevised regional stratigraphy. Journal of South Ameri-can Earth Sciences, 9, 153–160.

VON BLANCKENBURG, F. V., VILLA, I., BAUR, H.,MORTEANI, G. & STEIGER, R. H. 1989. Time cali-bration of a PT-path in the Western Tauern Window,Eastern Alps: the problem of closure temperatures.Contributions to Mineralogy and Petrology, 101,1–11.

WATERS, D. J. & MARTIN, H. N. 1993. Geobarometry inphengite-bearing eclogites. Terra Abstracts, 5,410–411; updated calibration of 1996 at World WideWeb Address: http://www.earth.ox.ac.uk/c.davewa/ecbar.html.

WAWRZYNIEC, T., MOLINA-GARZA, R. S., GEISSMAN,J. & IRIONDO, A. 2005. A newly discovered relic,transcurrent plate boundary – the Tonala shear zone

and paleomagnetic evolution of the western Mayablock, SW Mexico. Geological Society of AmericaAbstract with Programs, 37, 68.

WEBER, B., CAMERON, K. L., OSORIO, M. & SCHAAF, P.2005. A Late Permian tectonothermal event in Gren-ville crust of the Southern Maya terrane; U–Pbzircon ages from the Chiapas Massif, SoutheasternMexico. International Geology Review, 47, 509–529.

WEBER, B., IRIONDO, A., PREMO, W. R., HECHT, L. &SCHAAF, P. 2006. New insights into the history andorigin of the southern Maya Block, SE Mexico:U-Pb-SHRIMP zircon geochronology from meta-morphic rocks of the Chiapas Massif. InternationalJournal of Earth Sciences; doi: 10.1007/s00531-006-0093-7.

WEBER, B., VALENCIA, V. A., SCHAAF, P., POMPA-MERA, V. & RUIZ, J. 2008. Significance of prove-nance ages from the Chiapas Massif complex (SEMexico): redefining the Paleozoic basement of theMaya block and its evolution in a peri-Gondwananentourage. Journal of Geology, 6, 619–639.

WEILAND, T., SUAYAH, W. & FINCH, R. 1992. Petrologicand tectonic significance of Mesozoic volcanic rocks inthe Rıo Wampu area, eastern Honduras. Journal ofSouth American Earth Sciences, 6, 309–325.

WILSON, H.H. 1974. Cretaceous sedimentationand orogenyin nuclear Central America. American Association ofPetroleum Geologists Bulletin, 58, 1348–1396.

YANEZ, P., RUIZ, J., PATCHETT, P. J., ORTEGA-GUTIERREZ, F. & GEHRELS, G. E. 1991. Isotopicstudies of the Acatlan complex, southern Mexico:implications for Paleozoic North American tectonics.Geological Society of America Bulletin, 103, 817–828.


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The North American–Caribbean Plate boundary in Mexico ... · ocean-basin deposits (North El Tambor complex and El Pilar Formation, Fig. 2a) and crystalline basement of the Chuacu´s