ORIGINAL PAPER

Age constraints on Late Paleozoic evolution of continental crustfrom electron microprobe dating of monazite in the PeloritaniMountains (southern Italy): another example of resettingof monazite ages in high-grade rocks

Peter Appel • Rosolino Cirrincione •

Patrizia Fiannacca • Antonino Pezzino

Received: 27 October 2008 / Accepted: 18 December 2009 / Published online: 2 February 2010

� Springer-Verlag 2010

Abstract In situ monazite microprobe dating has been

performed, for the first time, on trondhjemite and

amphibolite facies metasediments from the Peloritani

Mountains in order to obtain information about the age of

metamorphism and intrusive magmatism within this still

poorly known sector of the Hercynian Belt. All samples

show single-stage monazite growth of Hercynian age. One

migmatite and one biotitic paragneiss yielded monazite

ages of 311 ± 4 and 298 ± 6 Ma, respectively. These ages

fit with previous age determinations in similar rocks from

southern Calabria, indicating a thermal metamorphic peak

at about 300 Ma, at the same time as widespread granitoid

magmatism. The older of the two ages might represent a

slightly earlier event, possibly associated with the

emplacement of an adjacent trondhjemite pluton, previ-

ously dated by SHRIMP at 314 Ma. No evidence for pre-

Hercynian events and only a little indication for some

monazite crystallization starting from ca. 360 Ma were

obtained from monazite dating of the metasediments,

suggesting either a single-stage metamorphic evolution or a

significant resetting of the monazite isotope system during

the main Hercynian event (ca. 300 Ma). Rare monazite

from a trondhjemite sample yields evidence for a late-

Hercynian age of about 275 Ma. This age is interpreted as

representing a post-magmatic stage of metasomatic mon-

azite crystallization, which significantly postdates the

emplacement of the original magmatic body.

Keywords Peloritani Mountains � Sicily �Monazite microprobe dating � Metapelite � Trondhjemite

Introduction

In the last decades the crustal evolution of the Calabria–

Peloritani Orogen has been the subject of considerable

work and Hercynian metamorphism and late-Hercynian

magmatism of both southern and northern sectors of the

orogen have started to be constrained on geochronological

grounds (Schenk 1980, 1990; Graeßner et al. 2000). Nev-

ertheless, for the southernmost termination of the Calabria–

Peloritani Orogen (Peloritani Mountains, north-eastern

Sicily), ages for magmatism and/or metamorphism are still

very scarce and mostly based on Rb/Sr and Ar–Ar systems.

A single high-precision in situ U–Pb zircon age of

314 ± 3 Ma was recently obtained for a trondhjemitic

sample (Fiannacca et al. 2008). A complicated structure

and a polyphase tectono-metamorphic history of the Pelo-

ritani Mountains, involving Alpine, Hercynian and possibly

pre-Hercynian events is postulated by most authors

studying this area (e.g. Atzori and Ferla 1992; Cirrincione

and Pezzino 1994; Ferla 2000 and references therein). In

such a context, in situ chemical dating of monazite appears

as a complementary method to obtain reliable constraints

useful to unravel the pre-Alpine history of the Peloritani

Mountains. Electron microprobe dating of monazite is a

non-destructive method, which offers the possibility of

dating monazite in its textural context allowing correlations

to be made between age data and discrete magmatic and

subsolidus events. The Aspromonte–Peloritani Unit, crop-

ping out in the north-eastern part of the Peloritani Moun-

tains and in the Aspromonte Massif of southern Calabria

(Fig. 1) was selected for this study since it stands out for its

P. Appel

Institut fur Geowissenschaften, Universitat Kiel,

24098 Kiel, Germany

R. Cirrincione � P. Fiannacca (&) � A. Pezzino

Dipartimento di Scienze Geologiche, Catania University,

Corso Italia 57, 95129 Catania, Italy

e-mail: pfianna@unict.it

123

Int J Earth Sci (Geol Rundsch) (2011) 100:107–123

DOI 10.1007/s00531-010-0511-8

widespread occurrence of trondhjemite plutons, which are

lacking in the other tectonic units of the Calabria Peloritani

Orogen. The present study reports the first monazite ages

obtained for two samples of high-grade metasedimentary

rocks and for a trondhjemite sample cropping out in the

Aspromonte–Peloritani Unit of the north-eastern Peloritani.

It is the aim of this study to contribute to a better under-

standing of the link between the tectono-metamorphic

evolution, including melt-generation and the post-mag-

matic processes, in this segment of the southern European

Hercynian Belt.

Geological setting

The Peloritani Mountain Belt (north-eastern Sicily) con-

stitutes the south-western termination of the Calabria–

Peloritani Orogen, a remnant of the Hercynian Belt,

which was reworked during the Alpine orogeny and that

now connects the Southern Apennine Chain and the

Maghrebid Chain (Fig. 1). The Peloritani Mountains

consist of a set of south-verging nappes of Hercynian base-

ment rocks, with metamorphic grade increasing towards

the top, and interposed fragments of Mesozoic–Cenozoic

Fig. 1 Geological sketch map

of Peloritani Mountains with

location of studied samples

(dark rectangles). a Structural

sketch map of the Peloritani

Mountains and the Aspromonte

Massif. b Basement-granitoids

relationships in the Calabria–

Peloritani Orogen. c Detailed

field map

108 Int J Earth Sci (Geol Rundsch) (2011) 100:107–123

123

sedimentary covers (Atzori and Vezzani 1974; Lentini and

Vezzani 1975).

The entire Peloritani Belt may be subdivided into two

complexes with different tectono-metamorphic histories

(Atzori et al. 1994; Cirrincione and Pezzino 1994). The

Lower Domain is exposed in the southern part of the

Peloritani Belt and comprises volcano-sedimentary Cam-

brian–Carboniferous sequences, which were affected by

Hercynian sub-greenschist to greenschist facies metamor-

phism and which now are covered by unmetamorphosed

Mesozoic–Cenozoic sediments. The Upper Domain, in the

north-eastern part of the belt, consists of two units (Man-

danici Unit and Aspromonte–Peloritani Unit), showing

Hercynian greenschist to amphibolite facies metamor-

phism, which in part are also affected by a younger Alpine

greenschist facies metamorphic overprint (Cirrincione and

Pezzino 1991; Atzori et al. 1994). Fragments of a meta-

morphosed Mesozoic–Cenozoic cover occur interposed

between the Mandanici Unit and the Aspromonte–Pelori-

tani Unit.

In the Peloritani Belt the Aspromonte–Peloritani Unit

represents the highest tectonic Unit, overlying the Man-

danici Unit, whereas in the Aspromonte Massif of southern

Calabria the same unit is sandwiched between the lower-

most Madonna di Polsi Unit (Pezzino et al. 2008) and the

uppermost Stilo Unit (Fig. 1). The most prominent rock

types in the Aspromonte–Peloritani Unit, in both northern

Sicily and southern Calabria, are middle crustal biotite

paragneisses and augen gneisses with minor amphibolites,

mica schists and marbles. The metamorphic rocks are

diffusely intruded by late-Hercynian weakly to strongly

peraluminous granitoids (D’Amico et al. 1982; Rottura

et al. 1990, 1993; Fiannacca et al. 2005a, 2008).

Late-Hercynian granitoids belong to two different

suites: a main (representing ca. 70% of the exposed rocks)

metaluminous to weakly peraluminous calc-alkaline

batholitic suite, and a strongly peraluminous suite, which is

composed of a number of small scale intrusions (Rottura

et al. 1993 and references therein). The granitoids are late-

to post-tectonic and were probably emplaced along ductile

shear zones connected to an extensional regime (Rottura

et al. 1990; Caggianelli et al. 2000, 2007). Only relatively

small plutons, mostly composed of weakly to strongly

peraluminous granitoids, are exposed within the medium-

high-grade basement of the Aspromonte–Peloritani Unit.

Among these granitoids, weakly peraluminous trondhje-

mites, not included in the frame of the late-Hercynian

magmatism by most of the previous authors, crop out either

as small plutons and stocks and as dm to m discordant to

sub-concordant leucosomes and dykes. Preliminary studies

on some of the trondhjemite bodies have suggested that

their origin is related to Hercynian crustal evolution pro-

cesses (Atzori et al. 1984a; Lo Giudice et al. 1985;

Fiannacca et al. 2005a). An origin by alkali metasomatism

at the expense of late-Hercynian granitoids was assumed

by Fiannacca et al. (2005a) for the Pizzo Bottino tron-

dhjemites, and a U–Pb zircon SHRIMP age of 314 ± 3 Ma

has been obtained for the emplacement of the original

magmatic body (Fiannacca et al. 2008).

Metamorphic conditions and previous

geochronological work

The timing of formation of the crystalline basement of the

Peloritani Mountains is still an open question. Most authors

ascribed it entirely to the Hercynian orogeny (e.g. Atzori

et al. 1984b; Ioppolo and Puglisi 1989; Messina et al.

1996), whereas others suggested that the Mandanici and

Aspromonte Units represent part of a pre-Hercynian

polymetamorphic basement (e.g. Ferla 1978, 2000; Bouillin

et al. 1987) or the result of welding of different pre-

Hercynian and Hercynian terranes during the final stages of

Hercynian orogeny (De Gregorio et al. 2003).

P–T estimates for the rocks of the Aspromonte–Pelori-

tani Unit in the Peloritani Mountains range between ca. 680

and 550�C at ca 5.0–3.0 kbar (Atzori et al. 1984b; Ioppolo

and Puglisi 1989; Messina et al. 1996; Rotolo and De Fazio

2001), with P–T peak conditions similar to those estimated

for the rocks of the Central Aspromonte Massif (650–

675�C at 4.0–5.0 kbar; Ortolano et al. 2005). Festa et al.

(2004, and references therein) additionally reported the

occurrence of pre/eo-Hercynian mafic granulites within

Hercynian migmatites of the northern Peloritani, with

Ca-rich Grt ? Cpx ? Qtz assemblages suggesting P &8–10 kbar and T & 700�C. A final widespread episode of

hydration under decreasing temperatures (&480�C) was

probably caused by the massive emplacement of metalu-

minous to strongly peraluminous late-Hercynian granitoids

at about 290 Ma (Rb–Sr data on micas; Rottura et al.

1990).

As for the pre-Hercynian evolution of southern Cala-

bria–Peloritani Orogen, different authors (Schenk and Todt

1989; Schenk 1990; Micheletti et al. 2007; Fiannacca et al.

2008, 2009) reported U–Pb zircon data from different

levels of the southern Calabrian and north-eastern Sicilian

crust indicating a late Pan-African/Cadomian (600–

500 Ma) crust-forming event. Micheletti et al. (2007)

obtained SIMS ages in the range of c. 560–525 Ma for the

magmatic protoliths of Calabrian augen gneisses, with

Archean (3.1 Ga), Paleoproterozoic (1.7–2.4 Ga) and

Neoproterozoic (0.6–0.9 Ga) inheritance. A similar set of

inherited ages and additional Pan-African inheritance has

been revealed by SHRIMP dating of zircon from a late-

Hercynian leucogranodiorite (Fiannacca et al. 2008). De

Gregorio et al. (2003) obtained a wide range of Ar–Ar

Int J Earth Sci (Geol Rundsch) (2011) 100:107–123 109

123

(stepwise heating method), Rb–Sr and U–Pb (stepwise

leaching method) ages for amphibole, biotite, muscovite,

K-feldspar and titanite, for samples of different rock types

from the north-eastern Peloritani, starting from about 1.6–

1.8 Ga up to 61–48 Ma. The latter ages are interpreted to

be related to Tertiary events localized along narrow shear

zones. The only pre-Hercynian ages are Ar–Ar hornblende

ages, from two amphibolite samples, interpreted as mixing

ages between a younger generation formed around 600 Ma

with older cores; U–Pb dating of one titanite from one of

those samples gave a still older age of 1.6–1.8 Ga.

Geochronological indications for early-Hercynian

events have been reported, from metapelites of southern

Calabria, by a Rb/Sr biotite age of ca. 330 Ma (Bonardi

et al. 1987), and by a poorly constrained lower concordia

intercept age of 377 ± 55 Ma (Schenk 1990). Neverthe-

less, more recently, Bonardi et al. (2008) indicated for the

same rocks Rb/Sr muscovite ages of ca. 314–308 Ma,

which are closer to recently obtained ages for the Hercy-

nian metamorphism.

In Sicily, De Gregorio et al. (2003) reported eo-Hercy-

nian 39Ar–40Ar hornblende ages ranging from 420 to

350 Ma and Hercynian 39Ar–40Ar hornblende and musco-

vite ages in the range of 340–300 Ma. Nevertheless, the

above authors interpret the first group of ages as the ages of

pre-Hercynian magmatic cumulates that escaped complete

Hercynian resetting, whereas no interpretation is provided

for most of the ages of the second group. Ar–Ar amphibole

ages of ca. 300 Ma were interpreted as the age of retro-

grade metamorphism of a former cpx–grt peak assemblage,

while a muscovite age of 301 ± 2 Ma from mylonitic

augen gneiss was interpreted to be related to nappe stack-

ing. The same authors obtained Ar–Ar biotite ages roughly

in the range ±240–50 Ma for the biotitic paragneisses from

the northern Peloritani. They reach the conclusion that the

basement of north-eastern Peloritani was mainly built in

the Paleozoic, but that pre-Hercynian relics and localized

Tertiary overprints also occur.

Atzori et al. (1990) indicated a common metamorphic

history for augen gneisses and associated biotitic parag-

neisses from the north-eastern Peloritani with Rb/Sr ages

on micas of 280–292 Ma, interpreted as cooling ages after

the Hercynian metamorphism.

U–Pb monazite ages (Graeßner et al. 2000) for similar

amphibolite facies paragneisses of the Aspromonte Massif

indicated a metamorphic peak at 295 to 293 ± 4 Ma (with

P–T conditions of 620�C at ca. 2.5 kbar for the base of the

upper crust), coeval with the lower crust (exposed in the

Serre Massif, southern Calabria), the latter characterized by

a peak temperature of 690–800�C at 5.5–7.5 kbar (Graeßner

et al. 2000 and reference therein). This metamorphic

peak was nearly synchronous with the bulk of granitoid

intrusions at 303–290 Ma (zircon, monazite and xenotime

U–Pb ages and whole-rock and mineral Rb–Sr ages; Borsi

and Dubois 1968; Borsi et al. 1976, Schenk 1980; Del

Moro et al. 1982; Graeßner et al. 2000; Fiannacca et al.

2008).

According to Graeßner and Schenk (1999) and Graeßner

et al. (2000) emplacement and crystallization of the large

volumes of granitoids were possibly responsible for a

regional scale late-stage metamorphism. This event

occurred under static conditions and resulted in a complete

recrystallisation of the mineral assemblages thus erasing

almost all records of previous tectonic and magmatic/

metamorphic stages, which are now only rarely preserved

(Caggianelli et al. 2007 and references therein).

Despite the absence of detailed geochronological

constraints, clockwise P–T–(t) paths inferred for the med-

ium- to high-grade rocks of the Aspromonte Massif and

Peloritani Mountains have been considered to be consistent

with processes of crustal thickening during early- and

middle-Hercynian collisional stages, followed by crustal

thinning, granitoid intrusion and unroofing during late-

Hercynian extensional stages (Festa et al. 2004; Caggia-

nelli et al. 2007). In particular, Atzori and Ferla (1992)

proposed distinct peaks for P and T in the northern Pelo-

ritani Mountains, with eo-Hercynian, or pre-Hercynian, P

peak (at ca. 600�C and 5.5 kbar) predating late-Hercynian

thermal peak (at ca. 630�C and 4.0 kbar).

Data reported for the Peloritani Mountains appear to be

consistent with a stage of low-P/high-T metamorphism at

about 300 Ma overprinting an older phase of Barrovian

metamorphism, as suggested for the whole orogenic seg-

ment (Graeßner and Schenk 1999; Graeßner et al. 2000;

Caggianelli and Prosser 2002; Festa et al. 2004).

Field relations and petrography

Amphibolite facies paragneisses and migmatites represent

the most common rock types in the Aspromonte–Peloritani

Unit. These rocks are the result of Hercynian metamor-

phism of greywacke sequences with variable pelitic com-

ponents (Lo Giudice et al. 1985, 1988). Their typical

mineral assemblage consists of quartz–plagioclase–biotite

with variable amounts of muscovite, K-feldspar, sillima-

nite, garnet, cordierite and andalusite. Synkinematic crys-

tallization appears to dominate in most paragneisses, but a

strong textural reorganization associated with crystalliza-

tion of plagioclase and micas is commonly evident, testi-

fying for post-deformational events.

Migmatites are mostly metatexites, with main foliation

defined by alternating millimetre- to decimetre-scale leu-

cosomes and biotite-rich meso-melanosomes. Leucosome

composition in the Peloritanian migmatites is leucogranitic

to trondhjemitic (Maccarrone et al. 1978; Atzori et al.

110 Int J Earth Sci (Geol Rundsch) (2011) 100:107–123

123

1985; Fiannacca et al. 2005b). Leucosomes studied by

Fiannacca et al. (2005b) have a trondhjemitic composition,

similar to melts produced by H2O-fluxed melting of mus-

covitic schists (Patino Douce and Harris 1998). However,

from petrographic investigations such trondhjemitic com-

positions often result to be a kind of ‘‘average’’ composi-

tion, which derives from the mixing of different

compositional domains. Similar ‘‘mixed’’ leucosomes,

suggesting multi-stage melting events, are reported for the

Serre metapelitic migmatites (Fornelli et al. 2002). The

mineralogical composition of Peloritanian leucosomes

appears further complicated by the presence of little

amounts of restite phases and the later occurrence of melt–

solid phases back-reactions and subsolidus alteration

(Fiannacca et al. 2005b).

Trondhjemitic rocks constitute variously sized scattered

bodies displaying a variety of field relations with the sur-

rounding metamorphic rocks (Fiannacca et al. 2005a and

references therein). The largest bodies are about 10 km2 in

extension and their contacts with the country rocks vary

from sharp (either discordant or concordant) to gradual.

Trondhjemites are sometimes associated with other leuco-

cratic rocks of granodioritic–monzogranitic compositions,

which have the same appearance in the field. Trondhje-

mites from the Peloritani Mountains are leucocratic and

mostly coarse- to very coarse-grained heterogranular rocks.

Variously sized metasedimentary enclaves of restitic/

xenolithic significance are frequently observed. The texture

is hypidiomorphic to autoallotriomorphic but a strong

subsolidus crystallization, related to metasomatism and

retrogression coupled with common solid-state textural

readjustment at depth, overprints the original magmatic

features, which are preserved within domains of variable

size.

Samples

Sample GC13 was collected close to a small trondhjemitic

body (Fig. 1). It is a fine-grained paragneiss composed of

quartz, plagioclase, biotite and very scarce K-feldspar.

Accessory phases are ilmenite, zircon, monazite and

unusually abundant apatite. Muscovite and chlorite only

occur as retrograde phases, grown at the expense of pre-

viously formed biotite. The sample displays a main folia-

tion S1 marked by synkinematic growth of biotite (Bt1),

plagioclase (Pl1), quartz, ilmenite and apatite. The S1

schistosity is partly affected by a coarse crenulation, which

does not lead to the formation of a new schistosity. A

second generation of coarser biotite (Bt2) cuts randomly

the foliation suggesting postkinematic growth. Thermo-

barometric data for similar paragneisses from the same area

suggest similar T conditions for the synkinematic and the

postkinematic stages, developed at T of ca. 550 and 550–

500�C, respectively. P conditions, estimated on the basis of

the phengite content of muscovite coexisting with K-feld-

spar (Massonne and Schreyer 1987) and of the Ca distri-

bution between garnet–plagioclase pairs (Ghent et al. 1979,

range from ca. 3.6 kbar, estimated for the synkinematic

stage, to 3.4–3.0 kbar obtained for the postkinematic one

(Ioppolo and Puglisi 1989).

Sample PB14 is from a metapelitic migmatite cropping

out at the southern side of the Pizzo Bottino trondhjemite

body (Fig. 1), which represents one of the largest tron-

dhjemite occurrences in the Peloritanian area (Fiannacca

et al. 2005a, b). The leucosomes (millimetre- to decimetre-

sized) consist of the typical mineral assemblage quartz–

plagioclase–muscovite–K-feldspar–sillimanite with minor

biotite and apatite and rare zircon and monazite as acces-

sories. They are always concordant within the foliation and

occur as lenticular bodies and as discontinuous layers,

which are often rimmed by millimetre-sized melanosomes.

Grain size of leucosomes is mainly coarse to very coarse.

PB14 leucosome-forming minerals are quartz, plagioclase,

muscovite, K-feldspar, sillimanite and small amounts of

biotite; accessory phases are apatite, zircon and rare

monazite. The overall composition of the leucosomes is

trondhjemitic with K-feldspar occurring as scattered

inclusions in plagioclase crystals or, less frequently, as an

interstitial and subhedral phase. Evidence for muscovite

breakdown under anhydrous conditions according to the

reaction muscovite ? quartz = K-feldspar ? sillima-

nite ? melt is locally given by textural relationships and

point to temperatures in excess of ca. 650�C, at pressure

greater than 3.5 kbar, as inferred for muscovite dehydra-

tion melting in the stability field of sillimanite. Widespread

growth of secondary plagioclase plus myrmekites and

muscovite–quartz symplectites at the expense of magmatic

microcline testifies the occurrence of significant crystalli-

zation following the anatectic stage (Fiannacca et al.

2005b). Meso and melanosomes show medium to fine grain

sizes and grano-lepidoblastic textures with some portions

characterized by postkinematic blastesis of micas and

plagioclase and the development of a granoblastic polyg-

onal texture. Mesosomes forming minerals are quartz,

plagioclase, biotite, muscovite and very little amounts of

K-feldspar. Accessory phases are apatite, zircon, monazite,

Fe–Ti oxides. Biotite is synkinematic, in single plats or in

lepidoblastic associations or, more rarely, postkinematic in

single plats. Muscovite is mostly synkinematic. Melano-

somes consist of millimetre-sized mafic selvages rimming

leucosomes and are mainly composed of biotite with lim-

ited amounts of sillimanite and muscovite. T estimates

available for similar Peloritanian migmatites (biotite–gar-

net and two-feldspar geothermometry; Ioppolo and Puglisi

1989) indicate T of ca. 640�C and 650–600�C for the

Int J Earth Sci (Geol Rundsch) (2011) 100:107–123 111

123

synkinematic and the postkinematic stages, respectively;

the postkinematic stage appears to have evolved at

decreasing pressures, in the range of 4.0–3.2 kbar.

Sample GC30 is a sample of a trondhjemite collected

from the Pizzo Bottino rock body (Fiannacca et al. 2005a;

Fig. 1). The selected sample is mainly composed of pla-

gioclase and quartz (ca. 90 vol%), and contains small

amounts of biotite, muscovite and K-feldspar. Accessory

phases are apatite, zircon, monazite and Fe–Ti oxides. The

sample is characterized by the occurrence of different

plagioclase populations represented by prevailing anhedral

to subhedral megacrysts and less abundant millimetre-sized

crystals. Several of the latter plagioclase crystals display

magmatic features, such as euhedral elongated habit and

simple twinning. Quartz forms medium to large anhedral

discrete grains or glomerocrystic aggregates; anhedral or

rounded quartz also occurs within the plagioclase. Biotite

and muscovite mainly occur as euhedral plates of variable

size with frequently corroded or fringed rims within the

plagioclase. Microcline occurs as scattered inclusions in

large plagioclase and in very rare interstitial patches.

Monazite age dating

Analytical techniques

All monazite analyses were performed with a JEOL JXA

8900 microprobe at the University of Kiel, equipped with

five WD spectrometers. For each analytical point a full

monazite analysis, consisting of 15 elements, was per-

formed. An accelerating voltage of 20 kV, a probe current

of 80 nA and a focused beam were used for all monazite

analyses. The JEOL H-type spectrometer with reduced

Rowland circle for high count rates was used for mea-

surements of lead, thorium and uranium. Background off-

sets were selected after long time fine WD scans of natural

monazite. The interference of Th Mc on U Mb was cor-

rected with an experimentally determined correction factor.

Counting times were adapted to net intensities to achieve

the desired objective of a low error for counting statistics at

reasonable counting times. Typical counting times for Pb

are 240 s on the peak and background. Under these con-

ditions and using natural crocoite as calibration standard,

the theoretical 1-sigma detection limit for Pb, as calculated

from the counting statistics of the background, is in the

range of 40–60 ppm. As standard materials, synthetic REE

orthophosphates (Jarosewich and Boatner 1991) corrected

for their Pb contents (Donovan et al. 2003) were used for P,

REE and Y; synthetic U-bearing glass for U; natural wol-

lastonite for Ca and Si; natural thorianite for Th; natural

crocoite for Pb; and corundum for Al. The JEOL ZAF

program was used for matrix correction of the monazite

analyses.

All monazite grains were measured in situ in petro-

graphic thin sections. A summary of the electron micro-

probe data for the analysed samples is given in Table 1.

Special polishing methods were used to avoid contamina-

tion by lead during sample preparation. To control the

quality of the data the internal laboratory standard F6 from

Manangoutry Pass, SE Madagascar (kindly provided by

Michael Raith, Bonn), was repetitively analysed during the

measurements. This monazite is homogeneous in compo-

sition and was dated with the U–Pb method of cogenetic

zircon and by a Sm–Nd monazite–biotite–garnet–zircon

isochron at 545 ± 2 and 542 ± 11 Ma, respectively

(Paquette et al. 1994), and recently by A. Moller (pers.

comm.) with TIMS at an age of 560 ± 1 Ma. Data with

significant Al contents were rejected to eliminate analyses

with secondary fluorescence artefacts that may occur if

analytical points are close to grain boundaries.

The analytical error for Th, U and Pb for each point was

calculated from counting statistics of the unknown and the

standard, using own software (for download at

http://www.ifg.uni-kiel.de/213.html).

After each monazite measurement, the apparent age

was calculated by solving the decay equation for the

U–Th–Pb system iteratively. All calculated apparent ages

of each sample were then evaluated to identify analyses

that clearly fall outside the range defined by a bell-shaped

normal distribution. This approach, together with a check

of the petrographic position of each measurement point,

helps to identify mixing ages without geological

significance.

Two methods were used to calculate ages for a set of

data: (1) the widely used chemical isochron method

(CHIME, Suzuki et al. 1991) and (2) the calculation of a

weighted average, which is frequently used in more recent

studies on chemical dating of monazite (e.g. Pyle et al.

2005). Isochron calculations were performed with the

CHIME-program of Kato et al. (1999). The isochron plot

has the advantage that it easily enables the identification

of data sets with distinct ages and additionally gives

information about the chemical variation of the analysed

monazite in terms of ThO2* and PbO. A drawback of this

method is that it typically suffers from high errors,

especially in cases when the variation of ThO2* is small,

and thus leads to a poorly defined regression line.

In contrast to this, the error of the weighted mean does

not depend on the chemical variation of the data and typ-

ically yields errors that are significantly lower than the

error calculated from the isochron method.

112 Int J Earth Sci (Geol Rundsch) (2011) 100:107–123

123

Table 1 Representative electron microprobe analysis and apparent ages of monazite from studied samples

Sample no. GC13 PB14 GC30

82 102 117 61 75 94 32 33 35

P2O5 31.10 30.91 30.65 30.71 26.39 30.82 30.66 30.58 30.83

SiO2 0.24 0.33 0.32 0.13 2.23 0.07 1.00 0.31 0.24

CaO 0.64 0.81 0.96 0.92 2.05 1.17 3.11 3.13 3.21

La2O3 15.28 14.36 13.83 14.44 10.22 12.93 8.21 8.30 8.05

Ce2O3 28.98 28.27 27.76 28.99 21.63 27.56 20.42 20.69 20.10

Pr2O3 3.09 3.02 3.05 2.97 2.41 2.98 2.45 2.47 2.52

Nd2O3 11.91 12.17 11.89 11.39 9.24 11.67 9.19 9.26 9.35

Sm2O3 1.90 2.10 1.99 1.69 1.71 2.05 3.52 3.60 3.65

Eu2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.15 0.21

Gd2O3 1.61 1.67 1.69 1.38 1.40 1.70 2.68 2.81 2.95

Dy2O3 0.51 0.49 0.60 0.41 0.64 0.64 1.03 1.06 1.15

Er2O3 0.13 0.05 0.06 0.15 0.15 0.14 0.08 0.08 0.10

Y2O3 1.97 1.90 2.04 2.18 2.17 2.39 2.79 2.98 3.07

ThO2 3.04 4.24 4.78 3.86 17.70 4.32 5.25 5.10 5.23

UO2 0.40 0.28 0.60 0.60 0.84 1.17 9.40 9.78 9.76

PbO 0.06 0.06 0.08 0.07 0.29 0.10 0.43 0.44 0.43

Al2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.43 0.05 0.02

Total 100.86 100.66 100.31 99.90 99.07 99.68 100.63 100.77 100.85

P 1.006 1.003 1.000 1.005 0.907 1.009 0.983 0.995 1.000

Si 0.009 0.013 0.012 0.005 0.090 0.003 0.038 0.012 0.009

Ca 0.026 0.033 0.040 0.038 0.089 0.049 0.126 0.129 0.132

La 0.215 0.203 0.197 0.206 0.153 0.184 0.115 0.118 0.114

Ce 0.405 0.397 0.392 0.410 0.322 0.390 0.283 0.291 0.282

Pr 0.043 0.042 0.043 0.042 0.036 0.042 0.034 0.035 0.035

Nd 0.163 0.167 0.164 0.157 0.134 0.161 0.124 0.127 0.128

Sm 0.025 0.028 0.027 0.023 0.024 0.027 0.046 0.048 0.048

Eu 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.002 0.003

Gd 0.020 0.021 0.022 0.018 0.019 0.022 0.034 0.036 0.037

Dy 0.006 0.006 0.007 0.005 0.008 0.008 0.013 0.013 0.014

Er 0.002 0.001 0.001 0.002 0.002 0.002 0.001 0.001 0.001

Y 0.040 0.039 0.042 0.045 0.047 0.049 0.056 0.061 0.063

Th 0.027 0.037 0.042 0.034 0.164 0.038 0.045 0.045 0.046

U 0.003 0.002 0.005 0.005 0.008 0.010 0.079 0.084 0.083

Pb 0.001 0.001 0.001 0.001 0.003 0.001 0.004 0.005 0.005

Al 0.000 0.000 0.000 0.000 0.000 0.000 0.019 0.002 0.001

Total 1.992 1.992 1.994 1.995 2.005 1.994 2.001 2.001 2.000

Xhu 0.004 0.007 0.008 0.002 0.085 0.001 0.003 0.004 0.001

Xmnz 0.872 0.857 0.838 0.850 0.663 0.819 0.627 0.624 0.615

Xxe 0.070 0.068 0.073 0.071 0.076 0.082 0.108 0.112 0.117

Xch 0.054 0.068 0.081 0.078 0.177 0.099 0.263 0.260 0.267

App. age (Ma) 309.4 271 274.6 298 338 300 284.8 280.9 277.5

Err (2-sigma) 46.3 38.4 30.2 36 12 25 9.1 8.5 8.1

Cations calculated on the basis of 4 oxygen

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Chemical composition and ages of monazite

Paragneiss GC13

Monazite of sample GC13 is generally smaller than

100 lm and anhedral with irregular-shaped grains

(Fig. 2a–c). Almost all grains are rich in small rounded

inclusions and have cracks, along which alteration has

affected the monazite. Some monazites are dismembered in

their outer parts. No significant zonation was observed in

backscattered electron images (BSEI). Monazite in this

sample either occurs in close association with biotite or

within the matrix. The first type of monazite is grown along

the grain boundaries of biotite or within biotite along

cracks. Matrix monazites are mostly smaller and more rare.

The chemical variation of monazite of GC 13 is

relatively restricted. According to the nomenclature of

Linthout (2007) all plot within the compositional field of

monazite (Fig. 3a). Th and LREE contents vary systemat-

ically, but to a small extent (Fig. 6a). U and Th do not

show any correlation (Fig. 4) and the plot of Y versus

HREE shows an unsystematic scatter of data points in a

wide range of the diagram (Fig. 6b).

In the Th ? U ? Si versus REE ? Y ? P diagram

(Fig. 3b) the analyses plot along a straight line close to the

cheralite exchange vector, thus indicating a single-stage

growth of monazite and the incorporation of Th and U into

the monazite structure mainly due to the cheralite

exchange.

The individual ages of these analyses scatter in a wide

range from 143 to 361 Ma. Plotted in a histogram (Fig. 7a),

they show a bimodal distribution with one maximum at

about 295 Ma and a second maximum at approximately

220 Ma. To calculate an age for the early phase of mon-

azite growth, analyses with apparent ages younger than

250 Ma cannot clearly be assigned to the younger or older

Gauss curve. All analyses with ages less than 250 Ma were

therefore excluded from the calculations of the high age.

The isochrone of sample GC13 thus is based on 38 anal-

yses and yields an age of 302 ± 58 Ma (2-sigma)

(Fig. 8a). The weighted mean age of 298 ± 6 (2-sigma)

Ma for these analyses is almost identical to the isochrone

age.

The reason for the large scatter with a bimodal age

distribution of the individual ages for sample GC13 first

Fig. 2 Backscattered electron images of representative monazite

grains in paragneiss (a–c sample GC13), migmatite (d–f sample

PB14) and trondhjemite (g–i sample GC30) from the Aspromonte–

Peloritani Unit of north-eastern Peloritani

Fig. 3 a Nomenclature diagram for the system 2REEPO4–

CaTh(PO4)2–2ThSiO4 after Linthout (2007). Analyses of all inves-

tigated samples fall within the compositional range of normal

monazite. Monazite of sample GC30 has a higher cheralite compo-

nent and also some patchy domains in monazite of sample PB14 have

significant huttonite and cheralite component. b Plot of

REE ? Y ? P versus Th ? U ? Si concentrations of monazites

from studied samples

114 Int J Earth Sci (Geol Rundsch) (2011) 100:107–123

123

appears to be unclear. There is no general pattern within

the age data which corresponds to the textural position of

the analyses. Moreover, within individual grains the ages

scatter but most of the younger ages occur in the more

rimward parts of the grains. To deduce the significance of

the younger ages of this sample, we calculated in a first

step a theoretical PbO content for each data point, which is

calculated from the weighted mean age for GC13 (298 Ma)

and from the measured ThO2 and UO2 concentrations. If

the Pb in this sample is undisturbed and not affected by

recrystallization and/or diffusion, the difference of this

value and the measured PbO has to correlate with the

difference between the calculated age of the analyses and

the mean age of the sample (Fig. 5a). Further, if a second

stage of monazite growth younger than that of the used

mean age (298 Ma) had occurred, then the respective data

points should be shifted towards DPbO values above the

linear data arrangement that is defined by data that reflect

an age of 298 Ma. Figure 5a shows a very good correlation

for most data points, and only few data points with low

apparent ages plot above the linear data arrangement. Y

contents of sample GC13 scatter in a wider range than in

the other samples and show only a very weak dependence

on the age (Fig. 5b).

The unsystematic behaviour of Y and HREE suggests

that monazite in this sample underwent disturbance in the

distribution of these elements after crystallization of the

monazite to different degrees. This, however, cannot be

applied to Pb, Th and U because the systematic correlation

of PbO to the reference of the mean age of 298 Ma as

displayed in Fig. 5a indicates that these elements are not

significantly affected by such element re-distribution.

Furthermore, the overall linear data arrangement in Fig. 5a

invalidates the interpretation that the young ages in the age

histogram (Fig. 7a) are geologically significant. It is thus

believed that monazite in GC13 crystallized during a sin-

gle-stage event at ca. 298 Ma.

Fig. 4 Plot of UO2 versus ThO2 concentrations. No systematic

correlation between these elements can be observed. Sample GC30

contains monazite with high U-rich content and sample PB14 shows a

wide range of ThO2 concentrations

-0.03

-0.02

-0.01

0

0.01

0.02

0.03

-300 -200 -100 0 100 200 300

Δ P

bO (

wt.-

%)

Δ Age (Ma)

GC13PB14GC30

(a)

(b)

0.01

0.02

0.03

0.04

0.05

0.06

Y (

c.p.

f.u.)

0.04

-0.04

Fig. 5 a DPbO versus Dchemical age diagram. DPbO is the

difference between a hypothetical PbO content, calculated from the

measured ThO2, UO2 and the weighted mean age of the sample and

the measured amount of PbO. Dchemical age is simply the difference

between the apparent age calculated for each analyses and the

weighted mean age for the sample. In case that the element

distribution of PbO, ThO2 and UO2 is undisturbed and the used

mean age is correct, all data points should plot along a straight line,

which crosses the origin of the axis. In case that a younger population

of data exists these data points should be shifted to higher DPbO.

b Chemical variation of Y (cations per 4 oxygen) versus Dchemical

age. The plot displays a weak tendency for low apparent ages of

analyses with low Y

Int J Earth Sci (Geol Rundsch) (2011) 100:107–123 115

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Migmatite PB14

The metapelitic sample PB14 contains abundant monazite.

Most of it occurs within the meso-melanosome, whereas

monazite in the leucosome is very rare. Monazite is closely

intergrown with biotite or can be found in the matrix close

to the grain boundaries with biotite. The grains are variable

in size, ranging from subhedral larger ones (50–100 lm) to

small anhedral grains which are mostly rounded (Fig. 2d–

f). Monazite is unzoned in BSEI except for some grains,

which include irregular-shaped bright patches that are

about 10–20 lm in size. These domains are enriched in

cheralite (up to 25 mol%) and huttonite (up to 17 mol%)

component with ThO2 contents up to 27 wt% and low

Ce2O3 values. Except from these, all analyses fall in the

compositional range of monazite (Fig. 3a) and have a rel-

atively restricted range of ThO2 between 3.5 and 6 wt%

(Fig. 4). Due to the overall large variation in ThO2 also the

LREE vary systematically, whereas Y and the HREE only

cover a restricted compositional range (Fig. 6a, b).

The apparent ages obtained from the analyses also vary

in a restricted range; 51 of a total of 60 analyses yield

apparent ages that are inside the range of a normal distri-

bution between 290 and 340 Ma. The isochrone age, based

on this set of analyses is 306 ± 20 Ma (2-sigma) and the

weighted mean age is 311 ± 4 (2-sigma) Ma (Figs. 7b,

8b). Again both ages are very close to each other. Also, the

ages of the Th-rich patches fall within the range of the

weighted mean age.

Trondhjemite GC 30

Only few grains of monazite were found in sample GC30.

They generally exhibit elongated shapes and are unzoned

(Fig. 2g–i). Most grains are very small (\10 lm) and

therefore only very few analyses could be obtained. All

show that monazite is enriched in cheralite component with

low huttonite component (Fig. 3a), and high content of

UO2 of more than 9 wt% (Fig. 4). Also, monazite from this

sample has the highest contents of Y and HREE which are

positively correlated (Fig. 6b). LREE contents of this

sample are relatively low (ca. 0.6 per 4 oxygen, Fig. 6a).

The ages of these analyses range between 250 and

290 Ma, when calculated from ThO2*, but age calculation

from UO2* yields similar results. The weighted mean age

is 275 ± 4 Ma (2-sigma). Because only five analyses could

be obtained from two tiny grains, no isochron was calcu-

lated for this sample. The ages represent only a rough and

provisional estimate for the time of monazite crystalliza-

tion and should be interpreted with caution. Nevertheless,

at the current state of knowledge they point to a somewhat

younger event compared to the ages of the semipelite and

the metapelite.

Discussion

Chemical dating with the electron microprobe was done on

monazite of two amphibolite facies metasedimentary rocks

and one trondhjemite from the basement of the Peloritani

Mountains. The data obtained by monazite dating of par-

agneiss and migmatite samples suggests a Hercynian evo-

lution for both lithotypes, with a metamorphic peak at

about 300 Ma. The rare monazite from the trondhjemitic

sample yields evidence for crystallization age of about

275 Ma.

In the interpretation of the in situ monazite ages

obtained by the present study it is also necessary to take

Fig. 6 Chemical variation diagrams for monazite analyses. a LREE

versus Th. These elements show a good correlation, which is mainly

caused by the huttonite exchange Th4? ? Si4? = LREE3? ? P5?.

b HREE versus Y plot. Y substitutes for REE and behaves similar to

HREE. The correlation between these elements is good for samples

PB14 and GC30 but for sample GC13 only a poor correlation can be

observed

116 Int J Earth Sci (Geol Rundsch) (2011) 100:107–123

123

into account the lack of pre-Hercynian ages in all the

analysed monazites, that would suggest an entirely

Hercynian evolution of the rocks of the Aspromonte–

Peloritani Unit. On the other hand, TIMS and ion-probe

U–Pb zircon Archean to Neoproterozoic–Early Cambrian

ages (Schenk 1990; Micheletti et al. 2007; Fiannacca et al.

2008, 2009) and U–Pb titanite and Ar–Ar hornblende

Proterozoic ages (De Gregorio et al. 2003) are reported for

magmatic and metamagmatic rocks of the same tectonic

unit in both southern Calabria and north-eastern Sicily.

By focusing first on the Hercynian evolution, the studied

unmigmatized paragneiss and the migmatite yielded mon-

azite ages of 298 ± 6 and 311 ± 4 Ma, respectively.

The age of 298 ± 6 Ma provides for the Peloritani

Mountains a scenario comparable to that already depicted

for the adjacent Aspromonte Massif of Southern Calabria,

where Graeßner et al. (2000) dated the Hercynian meta-

morphic peak at 295 to 293 ± 4 Ma based on ID-TIMS

U–Pb dating of monazite from similar amphibolite facies

paragneisses. Combined geochronological and petrological

data available for the whole southern Calabria–Peloritani

Orogen suggests that low-P/high-T metamorphism over-

printed the Barrovian one at about 300 Ma (Festa et al.

2004 and references therein). Graeßner and Schenk (1999)

and Caggianelli and Prosser (2002) suggested that this low-

pressure metamorphism was the result of combined

decompression and T increase due to widespread granitoid

plutonism. The above U–Pb ages were considered to rep-

resent the time of monazite growth and recrystallization

driven by the heat input provided by the large

Fig. 7 Weighted-histogram

representation of monazite age

data for the paragneiss (a) and

migmatite (b) samples. Each

small bell-shaped curve

represents the probability

density function for one

measurement. The dotted curveis the sum of all individual bell-

shaped curves. The thick curverepresents the weighted mean

age calculated by the statistical

procedure

Int J Earth Sci (Geol Rundsch) (2011) 100:107–123 117

123

metaluminous to strongly peraluminous granitoid intru-

sions into the middle crust (Graeßner et al. 2000). This last

Hercynian stage, occurring after the main deformational

events, obliterated the previous ones almost everywhere

and only rare relics allow constraining the prograde P–T

path (e.g. Acquafredda et al. 2006; Angı et al. 2010).

Evidence for mineral and textural readjustment con-

nected to thermal metamorphism and fluid–rock interaction

is widespread in the rocks of the north-eastern Peloritani. In

the studied samples post-tectonic crystallization of biotite

and plagioclase, associated with local foam texture devel-

opment or occurrence of feldspar replacement textures,

represent examples of such processes. The age of

311 ± 4 Ma of the migmatite sample, about 10 Myr older

than the age of the non-migmatitic paragneiss, might point

to a somewhat earlier event, since it precedes the bulk of

the Hercynian magmatism in the Calabria–Peloritani Oro-

gen. This event might be framed in a similar or in a post-

emplacement genetic context of the adjacent trondhjemite

body of Pizzo Bottino, which has a zircon SHRIMP mag-

matic crystallization age of 314 ± 3.5 Ma (Fiannacca et al.

2008).

Both obtained ages, however, have to be referred to a

thermal stage which postdated any blasto-deformational

event, since the 314 ± 3.5-Myr-old trondhjemite cut dis-

cordantly the migmatitic foliation, providing therefore

evidence for foliation development before 314 Ma.

Therefore, monazite ages obtained for metasedimentary

rocks from the Peloritanian segment of the Calabria–

Peloritani Orogen are in agreement with the scheme proposed

by Graeßner et al. (2000), explaining the metamorphic

peak recorded in southern Calabria at about 300 Ma as the

Fig. 8 PbO versus ThO2*

isochron diagram for the

paragneiss (a) and migmatite

(b) samples. The size of the

X- and Y-bars of each data point

represents the 2-sigma error,

calculated from the counting

statistics of the unknown and

the standard. Data points

marked with a square were not

used for the calculation of the

isochron and for the calculation

of a weighted mean age

118 Int J Earth Sci (Geol Rundsch) (2011) 100:107–123

123

result of combined decompression and thermal overprint

related to magmatism.

Based on available data, no clear evidence for an early-

Hercynian event related to the thickening stage or to the

early crustal melting processes leading to migmatite for-

mation may be derived from the present monazite dating

study. The data population of the apparent ages that range

up to 340 Ma from meso-melanosomes of the studied

migmatite (PB14) and up to 361 Ma in the non-migmatitic

paragneiss (GC13) may only indicate monazite crystalli-

zation during prograde metamorphism.

Trondhjemitic leucosomes produced by low-tempera-

ture H2O-present melting might be the best candidates to

represent the earliest appearance of Hercynian melt in the

Peloritani Mountains, and trondhjemitic leucosomes,

interpreted as produced under prograde metamorphism, are

reported from the lower crustal section of the Serre Massif

(Fornelli et al. 2002). Nevertheless, leucosome monazite

has not been dated because only one monazite grain was

found in the studied sample.

Different explanations can account for the monazite

ages obtained for the studied migmatite, by considering

that petrographic features and chemical data of these

migmatites indicate a complex history involving both fluid-

saturated and fluid-absent partial melting processes fol-

lowed by retrogression and metasomatism (Fiannacca et al.

2005b). In particular, H2O-present melting was probably

responsible for generations of quartz–plagioclase domi-

nated leucosomes, whereas fluid-absent melting generated

K-feldspar-rich granitic melt which locally appears to have

mixed to variable extent with the trondhjemitic leuco-

somes. Thus, the monazite age of 311 ± 4 Ma could be

related to the second event of partial melting experienced

by the rocks, leading to the formation of ‘‘mixed’’ tron-

dhjemite–granite leucosomes, as also reported for the Serre

migmatites (Fornelli et al. 2002). Another possibility,

supported by the widespread occurrence of feldspar

replacement textures, might invoke monazite dissolution

and recrystallisation directly caused by the infiltration of

fluids released from the adjacent crystallizing magmatic

Pizzo Bottino pluton. These interpretations are consistent

with the depicted tectono-metamorphic evolution of the

southern Calabria–Peloritani Orogen. Similar evidence

may be obtained by comparison with adjacent segments of

the southern Hercynian Belt, part of which was north-

eastern Sardinia before the post-Oligocene southeastern

drifting of the Calabria and Sicily basements to their

present position. In Sardinia, in situ Ar–Ar white mica ages

and U–Pb zircon ages of 350–320 Ma (Di Vincenzo et al.

2004; Palmeri et al. 2004; Giacomini et al. 2006) have been

obtained for the collision-related metamorphism. Meta-

texites with trondhjemitic leucosomes and Rb–Sr whole-

rock age of 344 ± 7 Ma (Ferrara et al. 1978) probably

represented the first occurrence of melt under prograde

metamorphism in the basement of north-eastern Sardinia

and, indeed, have been considered to reflect the collision

stage or the beginning of exhumation (Giacomini et al.

2006 and references therein). Lower amphibolite facies

mineral associations were produced during exhumation at

320–300 Ma, and finally, the last exhumation stages were

marked by granitoid plutonism at about 310–290 Ma,

involving high-temperature low-pressure metamorphic

overprints.

A similar tectono-metamorphic evolution, typical of

continental collision chains, with crustal thickening fol-

lowed by gravitative collapse, exhumation and granitoids

emplacement is reported from many other segments of the

European Hercynides such as Corsica (Menot and Orsini

1990), western and central Maures (Bellot et al. 2005);

French Massif central (Roig and Faure 2000; Costa and

Rey 1995; Rossi et al. 2006), and intra-Alpine massifs

(Paquette et al. 1989; von Raumer et al. 1999; Rubatto

et al. 2001).

The absence of pre-Hercynian ages, suggesting a solely

Hercynian single-stage metamorphic evolution, and of ages

reliably constraining the Hercynian collisional peak in the

studied samples might be ascribed to monazite recrystal-

lization. Similar situations are reported from many high-

grade rocks (e.g. Montel et al. 2000; Berger et al. 2005;

Braun and Appel 2006; Braun et al. 2007) where monazite

crystals totally included in garnet or quartz preserved older

ages compared to monazite crystals from the rock matrix or

from cleavage-bearing, fluid-accessible, mineral phases

such as orthopyroxene, kyanite or biotite. Communication

of monazite with other relevant Pb-containing phases in

metapelite (e.g. apatite, plagioclase) has been invoked as a

cause for monazite signature age resetting, independently

from further textural location or difference in grain size of

the monazite grains (Berger et al. 2005).

Dissolution/precipitation in the presence of a fluid phase

is also commonly considered a possible mechanism for

partial to complete resetting of the monazite isotope system

(Seydoux-Guillaume et al. 2002) and complete dissolution

and reprecipitation of monazite indeed has been frequently

ascribed to heat pulse and/or alteration by fluids released

by crystallizing magmatic bodies (e.g. Zeh et al. 2003 and

references therein).

In the studied metasedimentary rocks monazite occurs

as an inclusion in biotite or as a matrix phase, both types of

textural locations making monazite vulnerable to both

recrystallisation and fluid-induced dissolution/reprecipita-

tion leading to the resetting of the U–Pb system. Moreover,

although most of the youngest individual ages occur in the

outermost part of the grains, there is no systematic relation

between age data and the position of the analytical spots

within the monazite grains. Finally, diffuse alteration

Int J Earth Sci (Geol Rundsch) (2011) 100:107–123 119

123

affecting monazite along cracks may provide further sup-

port for monazite resetting late in the metamorphic history

of the rocks.

The monazite study of the trondhjemite sample did not

allow to constrain the emplacement age of the trondhjemite

pluton and to obtain clear evidence for the occurrences of

different crystallization stages of magmatic and metaso-

matic monazite, due to the very low amount and small size

of monazite in the sample analysed.

Nevertheless, these results provide preliminary infor-

mation about the age of monazite growth in the Pelorita-

nian trondhjemitic rocks and in conjunction with data

provided by recent studies (Fiannacca et al. 2005a, 2008),

they add new facts useful to gain understanding of the

nature and origin of these rocks.

First, the composition of trondhjemite monazite plots

in a very restricted field of the ternary diagram 2REE-

PO4–CaTh(PO4)2–2ThSiO4 (Fig. 3a), indicating a chera-

lite-enriched composition with low huttonite component.

Similar monazite, with cheralitic component greater than

that normally found in granites (6–18%, according to Bea

1996), has been mainly reported from, but is not

restricted to, strongly peraluminous S-type granites

(Forster 1998). Forster (1998) also reports that cheralite

appears to crystallize during late-stage processes from

fluid-rich residual liquids rather than during early mag-

matic crystallization. Cheralite-rich monazite is also

found as a secondary phase, for example associated to

chloritization affecting peraluminous granites (Poitrasson

et al. 1996).

Second, the preliminary chemical ages obtained from

dating of sample GC30 seem to indicate that monazite

crystallization occurred at about 275 Ma and, in any case,

not earlier than 290 Ma. This data strongly conflicts with

recent U–Pb SHRIMP dating, which reliably documents

crystallization of magmatic zircon at 314 ± 3.5 Ma in a

sample from the same trondhjemite body (Fiannacca et al.

2008). At present, the only possible explanation appears

that, if the chemical ages really reflect crystallization of

monazite at about 275 Ma, this crystallization should have

occurred under post-magmatic conditions, possibly during

the same metasomatic event invoked to account for the

unusual petrographical and geochemical features of the

trondhjemites (Fiannacca et al. 2005a). This interpretation

appears to be consistent with the cheralite-rich composition

of studied monazite.

Conclusion

The present study reports the first ‘‘in situ’’ chemical ages

of monazite from two metasedimentary rocks and one

trondhjemite from a former middle crustal sector of the

Calabria–Peloritani Orogen. The data obtained for one

biotitic paragneiss and one migmatite shows that they

shared the same Hercynian metamorphic evolution, cli-

maxing at ca. 300 Ma, in a low-pressure–high-temperature

event, which occurred in a context of combined decom-

pression and magmatism, a scenario which is generally

accepted for the entire Calabria–Peloritani Orogen at that

time. The age of 311 ± 4 Ma of the migmatite sample,

about 10 Ma older than that of the non-migmatitic parag-

neiss, might point to a somewhat earlier event. This might

be framed in the same genetic, or post-emplacement,

context of the adjacent trondhjemite body of Pizzo Bottino,

having a zircon SHRIMP magmatic crystallization age of

314 ± 3.5 Ma (Fiannacca et al. 2008).

No clear evidence has been found for both metasedi-

mentary samples, which allows temporal constraints on the

previous Barrovian metamorphic peak linked to the thick-

ening stage.

Some individual dates, however, are on the high age side

of the Gauss distribution and are either without any geo-

logical significance or represent relics of monazite crys-

tallization during prograde metamorphism. In either case,

the absence of monazite ages older than 361 Ma indicates a

solely Hercynian single-stage metamorphic evolution for

the metasedimentary rocks of the northern Peloritani

Mountains. Nevertheless, monazite occurrence in textural

locations favourable to isotopic resetting, namely as a

matrix phase or as an inclusion in cleavage-bearing mineral

phases, leaves open the possibility that the record of older

events could have been completely erased during the

Hercynian metamorphic evolution of the rocks.

The results of in situ monazite chemical dating of the

trondhjemite sample have to be considered preliminary at

present, due to the very low amount and small size of

analysed monazites. There results represent, however, the

first chemical and geochronological data of monazites from

trondhjemites of the Peloritani Mountains and, they

therefore, give a contribution to the knowledge of these

rocks. The chemical ages obtained in this work indicate the

starting of monazite crystallization after 290 Ma, with a

weighted mean age of 275 ± 4 Ma. Although this data has

to be considered as provisional, it would indicate a crys-

tallization age of the monazite too far ahead in time

compared to the real age of the magmatic crystallization, at

ca. 314 Ma, of the zircon from the same plutonic body, to

be considered a magmatic age. Consequently, monazite

ages obtained for the studied sample are better interpretable

as related to a post-magmatic stage of subsolidus crystal-

lization, in accordance with the metasomatic model pro-

posed by Fiannacca et al. (2005a) to explain the origin of

these peculiar trondhjemites.

120 Int J Earth Sci (Geol Rundsch) (2011) 100:107–123

123

Acknowledgments We thank A. Berger, an unknown reviewer and

the topic editor I. Braun for critical comments and suggestions that

helped significantly to improve the manuscript.

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