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. 2012 Mar;134-135(3-3):108–122. doi: 10.1016/j.lithos.2011.12.003

The behaviour of monazite from greenschist facies phyllites to anatectic gneisses: An example from the Chugach Metamorphic Complex, southern Alaska

Deta Gasser a,⁎, Emilie Bruand a,1, Daniela Rubatto b, Kurt Stüwe a
PMCID: PMC4599629  PMID: 26525358

Abstract

Monazite is a common accessory mineral in various metamorphic and magmatic rocks, and is widely used for U–Pb geochronology. However, linking monazite U–Pb ages with the PT evolution of the rock is not always straightforward. We investigated the behaviour of monazite in a metasedimentary sequence ranging from greenschist facies phyllites into upper amphibolites facies anatectic gneisses, which is exposed in the Eocene Chugach Metamorphic Complex of southern Alaska. We investigated textures, chemical compositions and U–Pb dates of monazite grains in samples of differing bulk rock composition and metamorphic grade, with particular focus on the relationship between monazite and other REE-bearing minerals such as allanite and xenotime. In the greenschist facies phyllites, detrital and metamorphic allanite is present, whereas monazite is absent. In lower amphibolites facies schists (~ 550–650 °C and ≥ 3.4 kbar), small, medium-Y monazite is wide-spread (Mnz1), indicating monazite growth prior and/or simultaneous with growth of garnet and andalusite. In anatectic gneisses, new low-Y, high-Th monazite (Mnz2) crystallised from partial melts, and a third, high-Y, low-Th monazite generation (Mnz3) formed during initial cooling and garnet resorption. U–Pb SHRIMP analysis of the second and third monazite generations yields ages of ~ 55–50 Ma. Monazite became unstable and was overgrown by allanite and/or allanite/epidote/apatite coronas within retrograde muscovite- and/or chlorite-bearing shear zones. This study documents polyphase, complex monazite growth and dissolution during a single, relatively short-lived metamorphic cycle.

Keywords: Monazite, Allanite, SHRIMP U–Pb geochronology, Alaska, Chugach Metamorphic Complex

Highlights

► Monazite/allanite stability from greenschist-facies phyllites to anatectic gneisses. ► Monazite-in reaction at 540 °C and ≤ 3.4 kbar. ► Dissolution and recrystallisation of monazite during partial melting, 54–51 Ma. ► Destabilisation of monazites related to fault zones.

1. Introduction

Monazite is a common accessory mineral in various metamorphic and magmatic rocks, particularly in rocks of felsic compositions, and it is an important LREE carrier (e.g. Corrie and Kohn, 2008, Spear, 2010, Spear and Pyle, 2002). Monazite incorporates relatively high concentrations of U and Th, has typically low concentrations of common Pb and is highly resistant to diffusive Pb-loss, and is thus well-suited for U–Pb geochronology (e.g. Harrison et al., 2002, Parrish, 1990). However, linking monazite ages with the PT evolution of the host rock is not always straightforward (e.g. Vance et al., 2003).

Monazite may occur as detrital grains below or at greenschist facies, and is stable as metamorphic monazite in rocks of appropriate bulk composition up to ultra-high temperature conditions (e.g. Goncalves et al., 2004, Spear and Pyle, 2002). Numerous studies have investigated monazite stability in metasedimentary metamorphic rocks with variable whole rock composition and metamorphic grade, and different monazite-producing reactions have been proposed (Budzyn et al., 2011, Fitzsimons et al., 2005, Foster et al., 2000, Foster et al., 2004, Gibson et al., 2004, Harlov and Hetherington, 2010, Janots et al., 2006, Janots et al., 2008, Janots et al., 2009, Kim et al., 2009, Kingsbury et al., 1993, Kohn and Malloy, 2004, Kohn et al., 2005, Rasmussen et al., 2006, Rubatto et al., 2001, Smith and Barreiro, 1990, Williams et al., 2007, Wing et al., 2003). These studies highlight the importance of understanding the stability of monazite and other LREE minerals (e.g. allanite, xenotime) within a particular metamorphic sequence before a meaningful interpretation of U–Pb monazite formation dates can be obtained.

In this contribution, we present the results of an investigation of monazite stability within a single-cycle, metasedimentary sequence evolving from upper greenschist facies phyllites to amphibolite facies schists into upper-amphibolite facies migmatites. The investigated sequence is exposed along the southern Alaskan margin within a Late Cretaceous to Paleocene accretionary prism and is termed the Chugach Metamorphic Complex (CMC, Hudson and Plafker, 1982, Sisson et al., 1989, Gasser et al., 2011). We describe the textural relationships of monazite in samples of differing bulk rock composition and metamorphic grade, with particular focus on the relationship of monazite to other REE-bearing minerals such as allanite and xenotime. The composition of monazite from the two metamorphic zones where monazite is stable is monitored, and U–Pb SHRIMP dates for selected monazite samples are used to constrain the age of different monazite growth events. Our study documents the complexity of monazite growth and recrystallisation during a single-cycle, relatively short-lived metamorphic event.

2. Geology of the Chugach Metamorphic Complex

The southern Alaskan margin is dominated by a large-scale, complexly deformed Cretaceous to Paleogene accretionary prism which stretches ~ 2200 km from Sanak Island in the west to Baranof Island in the east (the Chugach and Prince William terranes, Fig. 1a, Plafker et al., 1994). This accretionary prism mainly consists of turbiditic rocks: conglomerates, impure sandstones, marls and claystones, which are locally intercalated with mafic volcanic rocks (e.g. Bruand et al., 2011, Nielsen and Zuffa, 1982, Sample and Reid, 2003). The accretionary prism is mainly metamorphosed in phrenite–pumpellyite to greenschist facies, but it is metamorphosed to upper amphibolite facies within the Chugach Metamorphic Complex (CMC, Fig. 1b, Bruand et al., 2011, Gasser et al., 2011, Hudson and Plafker, 1982, Pavlis and Sisson, 1995, Sisson et al., 1989).

Fig. 1.

Fig. 1

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(a) Overview map of Alaska and Canada with the location of Fig. 1b indicated. (b) Geological overview map of the Chugach Metamorphic Complex (CMC) within the Chugach terrane of southern Alaska. Small inset maps show stippled glacier outlines and the localities of the investigated samples as well as PT conditions for the different metamorphic zones after Bruand (2011). Note the location of sample Loc4 outside the small inset maps. Abbreviation: BRF = Border Range Fault System.

The CMC itself consists of two macroscopically different metamorphic zones: an outer schist zone composed of fine-grained biotite–quartz–plagioclase schist, which surrounds an inner gneiss zone composed of layered migmatitic gneisses (Fig. 1, Hudson and Plafker, 1982). In the following, we additionally refer to the greenschist facies phyllites outside the CMC as the phyllite zone. The distribution of these zones is highly asymmetric: there is a wide transition from phyllites into gneisses in the north and an abrupt transition from gneisses to phyllites along ductile to brittle faults in the south (Fig. 1; Gasser et al., 2011). The mineralogy observed in the different zones varies depending on bulk rock composition and metamorphic grade. Bulk rock compositions vary from psammites to calcareous metapelites to metapelites (Bruand, 2011). In metapelitic rocks, andalusite is wide-spread in the western and northern schist zone of the CMC and cordierite occurs locally in the western gneiss zone. Sillimanite occurs in the southern schist zone and is abundant over the entire gneiss zone, but is absent in the highest-grade rocks of the gneiss zone (Bruand, 2011, Hudson and Plafker, 1982, Pavlis and Sisson, 1995, Pavlis and Sisson, 2003, Sisson and Pavlis, 1993, Sisson et al., 1989). Garnet is usually abundant in metapelitic to calcareous metapelitic rocks in both the schist and gneiss zones, whereas K-feldspar occurs rarely in migmatites of the gneiss zone (Bruand, 2011, Sisson et al., 1989). Metamorphosed mafic rocks exposed along the southern border of the CMC contain abundant amphibole and locally garnet (Sisson et al., 2003; Bruand et al., 2011). Sisson et al. (1989) and Bruand (2011) derived PT conditions of 400–550 °C and < 4 kbar for the phyllite zone, ~ 500–650 °C and ~ 3–7 kbar for the schist zone, and ~ 650–750 °C and 4–13 kbar for the gneiss zone, with the highest pressures occurring in the southern and eastern parts of the gneiss zone (Fig. 1, Table 1).

Table 1.

List of samples and their respective mineralogy and P–T estimates.

Sample Lat (°N) Long (°W) Metam. zone Mineralogy
+ Qz + Pl + Bt		 Mnz Mnz(i) Xtm Ap Ep/Aln Zrn Ilm Ttn Py Hem Rt T(°C)a p (kbar)a
B43 60.9263 143.2130 Phyll. Ms–Chl–Ep x x x x x x x 515 ± 50b
B44 60.9263 143.2130 Phyll. Ms–Chl–Ep x x x x x 535 ± 50b
B45 60.9328 143.2025 Phyll. Ms–Chl–Ep x x x x x x 547 ± 50b
T28 60.6704 142.9123 Phyll. Ms–Chl–Ep x x x x x x 432 ± 50b
Tk14 60.6941 141.5803 Phyll. Ms–Chl–Ep x x x x 536 ± 50b
Tk15 60.6965 141.5739 Phyll. Ms–Chl–Ep x x x x x 526 ± 50b
Tk12 60.6878 141.5807 Schist And–Ms–Grt x x x x x 543 ± 23c 3.4 ± 0.9c
Tk10 60.6820 141.5716 Schist And–Ms–Grt x x x x x x x 645 ± 49c 5.9 ± 2.1c
Tk9 60.6752 141.5608 Schist And–Ms–Grt x x x x x x x x > 640b
Tk8 60.6731 141.5605 Schist (And)Sil–Ms–Grt x x x x x x
Tk6 60.6642 141.5488 Schist (And)Sil–Ms x x x xr x x x
Tk3 60.6642 141.5488 Schist (And)Sil–Ms–Grt x x x x x x 658 ± 33c 6.4 ± 1.4c
B39 60.8858 143.2384 Schist Grt–Sil–Ms x x x x xp?,xr x x x 597 ± 50b
B40 60.8883 143.2384 Schist Grt–Sil–Ms x x x xr x x x
B9–2 60.8583 143.2362 Schist Grt–Sil–Ms x x x xr x x x 632 ± 30d 6.8 ± 1.6c
B22 60.8124 143.3148 Gneiss Grt–Sil x x x x x x 623 ± 30d 4.7 ± 1.9c
B23 60.8110 143.1928 Gneiss Grt–Sil x x x xr x x x x 645 ± 30d
B26 60.8236 143.2033 Gneiss Grt–Sil–Ms x x x x x x 628 ± 30d 5.6 ± 2c
KB1 60.7643 143.2881 Gneiss Grt–Ms x x x x x x 645 ± 30d 4.7 ± 1.6c
KB2 60.7643 143.2881 Gneiss Grt x x x xr x x x 666 ± 30d 5.1 ± 1.6c
KB5 60.7381 143.3015 Gneiss Ms x x x x x x
T27 60.7146 142.9147 Gneiss Grt–Ms x x x xr x x x 679 ± 32c 7.3 ± 1.1c
Loc4 60.5544 141.3030 Gneiss Grt–Sil–Ms x x x x x 666 ± 25c 8.2 ± 1.2c
B21 60.8124 143.3148 Gneiss x x x x x x
T7 60.6960 142.9031 Gneiss Grt–Amp–Ms–Chl xu x x x x
T15a 60.6946 142.9069 Gneiss Grt–Ms–Chl x x x x,xr x x x x 683 ± 33c 7.3 ± 1.1c
N7 59.8116 138.8131 Gneiss Grt–Ms–Chl xu x x x x x x
N17 59.8330 138.8322 Gneiss Grt–Ms–Chl xu x x x x x x
N38 59.7914 138.8207 Gneiss Amp–Grt xu x x x x x 625 ± 30d 8.3 ± 1.2c
N39 59.7914 138.8207 Gneiss Amp–Grt x x x x x x
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x = mineral is present in matrix, xu = mineral is unstable, Mnz(i) = monazite as inclusion in garnet and/or andalusite, xr = mineral occurs as retrograde fibres around monazite, xp = mineral is present as unstable prograde crystal. Mineral abbreviations according to Whitney and Evans, 2010

a

PT data from Bruand (2011).

b

Graphite thermometry.

c

Average PT from thermocalc.

d

Biotite–Garnet thermometry.

The development of the CMC occurred during a remarkably short time span. Detrital zircon within the metasediments indicates deposition after ~ 65–60 Ma (Gasser, 2010, Gasser et al., in press, Kochelek et al., 2011). Based on U–Pb SHRIMP dates of thin zircon overgrowths in the migmatites of the gneiss zone, peak metamorphic conditions were reached ~ 5 m.y. later at ~ 55–51 Ma virtually simultaneously at least ~ 300 km along-strike of the complex (Gasser et al., in press). This short metamorphic event was followed by rapid cooling through 40Ar/39Ar muscovite and biotite closure temperatures (~ 480–320 °C) at 50–45 Ma in the western and central parts of the complex and slower cooling through these closure temperatures at 25–15 Ma in the southeastern part of the complex (Bradley et al., 1993, Gasser et al., 2011, Hudson et al., 1977a, Hudson et al., 1977b, Hudson et al., 1979, Sisson et al., 1989; Sisson et al., 2003).

3. Analytical techniques

Major element whole rock compositions of all investigated samples were obtained with a Siemens Pioneer S4 WD-XRF at the Department of Earth Sciences, University of Graz (Austria). Identification and high-contrast back-scattered electron (BSE) imaging of accessory minerals was done on polished thin sections at the Department of Earth Sciences, University of Graz (Austria) with a JEOL JSM-6310 scanning electron microscope, equipped with an IDIX EDX system for mineral identification, working at 15 kV, ~ 6 nA and 15 mm working distance. Element mapping and quantitative chemical analyses of monazite grains were obtained on a JEOL JXA 8200 microprobe at the Department for Applied Geosciences, University of Leoben (Austria). X-ray maps of monazite grains were obtained with 30 kV, ~ 10 nA, step size of ~ 0.2–0.4 μm and dwell times of 20–50 ms. Quantitative analyses of monazite were obtained using synthetic and natural mineral standards and a ZAF data correction scheme. Accelerating voltage was ~ 20 kV and beam current was 50 nA. Kα X-ray lines were used for Si, P and Ca, Lα X-ray lines were used for Y, Ce, Dy, La, Pr, Nd, Eu, Ho, Tb and Er, Lβ X-ray lines were used for Sm and Gd and Mα X-ray lines were used for Th. U and Pb were not analysed.

Monazite grains were separated for U–Pb SHRIMP dating from fist-sized samples by crushing, sieving, magnetic and heavy liquid separation, and then handpicked, mounted in epoxy and polished to expose the grain centres. BSE imaging of the separated monazite grains was done at the Electron Microscope Unit, Australian National University (ANU), with a Cambridge S360 scanning electron microscope working at 15 kV, ~ 2 nA and a working distance of ~ 17 mm. U–Pb analysis on monazite grains was performed with a sensitive high-resolution ion microprobe (SHRIMP II) at the Research School of Earth Sciences (RSES), ANU, using a beam size of ~ 20 μm. The data were collected in sets of six scans throughout the masses, and reference monazite 44,069 (424.9 ± 0.4 Ma; Aleinikoff et al., 2006) was analysed each fourth analysis. Energy filtering was applied to eliminate the molecular interference on 204Pb (Rubatto et al., 2001) and to reduce any matrix effect. The analyses were corrected for common Pb based on 204Pb and 207Pb measurements following Williams (1998). The two correction methods deliver results indistinguishable within error. Th–Pb dates, despite being less precise, are generally in agreement with the U–Pb dates. Here, 204Pb corrected data are presented in Table 5 and plotted on concordia diagrams, and 206Pb/238U weighted mean dates were calculated. The common Pb composition was modelled according to Stacey and Kramers (1975). U–Pb data were collected over one analytical session using the same standard, with the session having a calibration error of 2.7% (2σ), which was propagated through to single analyses. Additionally, a 1% error was added to the final mean dates to account for external errors. Data evaluation and calculation of dates were done using the software Squid1 and Isoplot/Ex 3, respectively (Ludwig, 2003). Mineral abbreviations are according to Whitney and Evans (2010).

Table 5.

Monazite U–Pb shrimp data.

Label U (ppm) Th (ppm) 232Th/238U % Pb common 207Pb/235U 1σ % 206Pb/238U 1σ % error correlation 206Pb/238U Age (Ma) ± 1σ
Sample B21
B21-1C 4112 38,887 9.8 1.42 0.0474 8.2 0.00809 1.6 0.196 52.0 0.8
B21-2C 3196 26,546 8.6 0.00 0.0552 3.4 0.00784 1.6 0.469 50.4 0.8
B21-2M 5441 64,526 12.3 0.71 0.0480 5.1 0.00809 1.8 0.357 52.0 0.9
B21-2R 3080 37,389 12.5 0.37 0.0517 5.1 0.00833 1.6 0.318 53.5 0.9
B21-3C 3972 38,741 10.1 0.36 0.0497 3.9 0.00824 1.5 0.394 52.9 0.8
B21-4C 1603 23,122 14.9 1.16 0.0455 7.0 0.00795 1.8 0.253 51.0 0.9
B21-5C 2885 53,646 19.2 0.31 0.0498 5.8 0.00793 1.6 0.279 50.9 0.8
B21-6C 2626 28,878 11.4 1.03 0.0466 7.6 0.00799 2.3 0.306 51.3 1.2
B21-6R 2471 22,713 9.5 0.00 0.0545 3.9 0.00795 1.8 0.474 51.0 0.9
B21-7C 4336 39,452 9.4 0.50 0.0451 6.6 0.00787 1.7 0.250 50.6 0.8
B21-8Ca 4943 68,740 14.4 1.35 0.0453 10.9 0.00836 2.1 0.197 53.6 1.1
B21-9C 2240 55,502 25.6 1.04 0.0457 10.4 0.00792 1.9 0.183 50.9 1.0
B21-10C 3932 29,956 7.9 0.27 0.0531 6.8 0.00817 3.0 0.446 52.4 1.6



Sample KB1
KB1A-11R 13,417 30,504 2.3 0.18 0.0493 3.0 0.00795 1.5 0.486 51.0 0.7
KB1A-10R 6339 31,408 5.1 0.58 0.0500 4.9 0.00798 1.5 0.306 51.3 0.8
KB1A-12R 8339 32,679 4.0 0.19 0.0522 2.7 0.00813 1.5 0.567 52.2 0.8
KB1A-8C 7415 30,480 4.2 0.42 0.0511 4.0 0.00816 1.5 0.389 52.4 0.8
KB1A-6C 8198 30,578 3.9 0.00 0.0537 2.7 0.00818 1.5 0.544 52.5 0.8
KB1A-1C 5546 27,957 5.2 0.00 0.0596 5.5 0.00822 1.6 0.285 52.8 0.8
KB1A-7C 7316 32,106 4.5 0.00 0.0503 8.5 0.00827 1.5 0.177 53.1 0.8
KB1A-5C 6360 29,322 4.8 0.41 0.0527 3.4 0.00833 1.5 0.436 53.5 0.8
KB1A-3C 7383 27,492 3.8 0.42 0.0516 4.6 0.00835 2.0 0.426 53.6 1.0
KB1A-9C 2486 34,781 14.5 0.85 0.0563 9.4 0.00836 2.4 0.257 53.6 1.3
KB1A-4C 7866 28,833 3.8 0.49 0.0503 3.9 0.00838 1.5 0.374 53.8 0.8
KB1A-10C 3076 34,049 11.4 0.53 0.0509 6.6 0.00839 1.6 0.244 53.8 0.9
KB1A-2Ca 2023 21,587 11.0 4.12 0.0199 47.6 0.00705 3.0 0.063 45.3 1.4



Sample KB5
KB5-4C 3305 36,461 11.4 0.85 0.0479 5.6 0.00804 2.1 0.371 51.6 1.1
KB5-2C 2009 27,306 14.0 0.00 0.0577 4.4 0.00815 2.3 0.516 52.3 1.2
KB5-1C 6044 21,853 3.7 0.25 0.0560 3.7 0.00822 1.9 0.521 52.8 1.0
KB5-9C 5172 25,267 5.0 0.51 0.0496 4.3 0.00823 1.5 0.351 52.8 0.8
KB5-1OC 2688 30,811 11.8 0.00 0.0584 3.6 0.00830 1.8 0.508 53.3 1.0
KB5-1R 6867 22,129 3.3 0.00 0.0554 2.7 0.00830 1.9 0.675 53.3 1.0
KB5-8C 4537 41,541 9.5 1.13 0.0456 7.4 0.00831 1.5 0.208 53.4 0.8
KB5-7C 5012 22,550 4.6 0.52 0.0489 4.1 0.00836 1.6 0.395 53.7 0.9
KB5-1OR 10,058 18,487 1.9 0.00 0.0540 2.3 0.00838 1.4 0.635 53.8 0.8
KB5-5R 8488 20,678 2.5 0.28 0.0530 3.0 0.00843 1.8 0.590 54.1 1.0
KB5-6C 5221 38,398 7.6 0.60 0.0552 5.0 0.00844 1.6 0.314 54.2 0.8
KB5-3C 3660 31,100 8.8 0.00 0.0569 3.1 0.00845 1.6 0.499 54.2 0.8
KB5-4M 8347 20,185 2.5 0.00 0.0552 2.5 0.00853 1.6 0.652 54.8 0.9
KB5-4Ra 4402 26,956 6.3 1.18 0.0405 9.4 0.00784 3.7 0.391 50.3 1.9
KB5-8Ra 5323 21,729 4.2 4.42 0.0545 18.3 0.00802 1.9 0.103 51.5 1.0
KB5-2Ra 3941 19,976 5.2 2.38 0.0721 9.2 0.00827 1.6 0.178 53.1 0.9
KB5-5Ca 4375 37,643 8.9 1.41 0.0424 13.1 0.00828 1.6 0.124 53.1 0.9
KB5-8Ma 2528 29,537 12.1 1.10 0.0441 14.6 0.00851 3.0 0.202 54.6 1.6



Sample T27
T27-11C 3291 34,414 10.8 0.81 0.0497 6.8 0.00803 1.6 0.238 51.5 0.8
T27-7R 5519 31,414 5.9 0.59 0.0515 5.6 0.00808 1.6 0.280 51.8 0.8
T27-6C 4057 38,046 9.7 0.96 0.0504 6.0 0.00808 2.2 0.361 51.9 1.1
T27-10R 15,265 21,526 1.5 0.26 0.0506 3.1 0.00809 1.5 0.467 51.9 0.8
T27-1C 6439 68,741 11.0 1.13 0.0472 4.5 0.00810 1.7 0.364 52.0 0.9
T27-2R 8892 28,224 3.3 0.36 0.0489 3.1 0.00812 1.5 0.479 52.1 0.8
T27-4C 8999 22,585 2.6 0.33 0.0513 4.0 0.00814 2.9 0.709 52.2 1.5
T27-7C 3737 28,640 7.9 0.00 0.0586 3.2 0.00814 1.5 0.481 52.3 0.8
T27-9C 5246 31,195 6.1 0.80 0.0535 4.9 0.00818 2.3 0.464 52.5 1.2
T27-6R 8555 27,730 3.3 0.71 0.0492 3.9 0.00818 1.6 0.412 52.5 0.8
T27-3C 4515 36,312 8.3 0.00 0.0580 3.1 0.00829 1.6 0.507 53.2 0.8
T27-1R 9590 18,836 2.0 0.35 0.0474 4.7 0.00832 1.4 0.309 53.4 0.8
T27-8C 4677 44,039 9.7 0.46 0.0525 5.9 0.00833 2.1 0.356 53.5 1.1
T27-5C 11,855 27,660 2.4 0.38 0.0517 3.3 0.00838 1.9 0.572 53.8 1.0
T27-1Ma 1452 44,155 31.4 1.57 0.0465 14.2 0.00759 3.9 0.275 48.8 1.9
T27-2Ca 2796 32,899 12.2 1.50 0.0444 12.3 0.00791 2.2 0.183 50.8 1.1
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a

Denotes measurement excluded in age calculation and Concordia plots.

4. Results

In order to study the stability of monazite and other REE-bearing accessory minerals within samples of variable bulk rock composition and variable metamorphic grade across and along-strike of the Chugach Metamorphic Complex, we investigated a total of 30 samples (Fig. 1, Table 1). Sample numbers, coordinates, major and accessory minerals and PT conditions (Bruand, 2011) of all samples are summarised in Table 1. The major element bulk rock compositions of all samples are given in Table 2. Six samples are derived from the phyllite zone, nine samples from the schist zone, and the remaining 15 samples are from the gneiss zone of the Chugach Metamorphic Complex (Fig. 1, Table 1).

Table 2.

Major element bulk rock compositions of all samples in weight %.

Sample Rocka SiO2 (%) Al2O3 (%) Fe2O3 (%) MnO (%) MgO (%) CaO (%) Na2O (%) K2O (%) TiO2 (%) P2O5 (%) Sum (%)
B43 p 68.86 13.35 5.76 0.07 2.39 2.06 2.20 2.04 0.70 0.15 99.44
B44 p 61.07 15.63 7.08 0.12 3.79 3.33 2.75 2.22 0.91 0.18 99.29
B45 p 70.20 13.79 5.16 0.07 1.80 1.38 1.74 2.28 0.65 0.15 99.57
T28 p 59.88 17.28 6.67 0.13 3.03 2.89 2.66 2.52 0.76 0.19 99.58
Tk14 p 66.34 12.54 6.54 0.13 3.48 4.57 2.78 1.03 0.93 0.15 99.48
Tk15 p 60.78 16.63 6.79 0.12 3.00 5.07 4.42 0.89 0.77 0.18 99.77
Tk12 s 69.22 13.77 6.02 0.09 2.15 1.43 1.91 2.25 0.68 0.18 99.60
Tk10 s 64.36 16.32 7.34 0.10 2.60 1.16 1.47 2.88 0.84 0.24 99.72
Tk9 s 69.62 12.78 6.14 0.12 2.12 1.53 1.97 2.35 0.68 0.23 99.35
Tk8 s 66.50 16.04 5.44 0.07 2.04 1.86 3.67 2.29 0.69 0.06 100.01
Tk6 s 65.59 15.74 5.74 0.08 2.31 1.06 1.97 3.30 0.74 0.18 99.87
Tk3 s 59.95 18.57 7.82 0.10 2.90 1.22 2.25 3.17 0.97 0.20 99.91
B40 s 46.54 22.30 11.60 0.22 5.30 3.58 3.47 4.01 1.22 0.09 99.92
B39 s 60.26 17.43 7.26 0.23 2.90 3.16 2.81 3.06 0.75 0.23 99.57
B9 s 60.74 17.59 7.95 0.10 3.12 1.91 2.21 3.13 0.85 0.23 100.48
B21 g 68.08 15.17 4.48 0.07 1.97 3.30 3.28 1.99 0.59 0.17 100.01
B22 g 60.36 17.89 7.27 0.11 2.86 2.79 3.27 2.93 0.84 0.25 99.60
B23 g 60.84 17.16 6.95 0.10 2.87 2.80 3.09 2.73 0.76 0.28 98.83
B26 g 63.24 16.75 7.60 0.09 3.12 2.44 2.53 2.85 0.84 0.23 100.58
KB1f g 61.71 16.89 6.94 0.11 2.87 3.15 3.08 2.61 0.87 0.22 100.31
KB2 g 62.83 16.55 6.44 0.09 2.72 3.19 3.24 2.41 0.73 0.23 99.79
KB5rest g 62.34 17.14 6.31 0.11 2.84 3.30 3.20 2.93 0.86 0.18 100.70
T7 mv 56.53 17.81 11.49 0.23 4.41 4.49 2.57 0.27 1.07 0.24 100.95
T15a g 68.64 15.11 4.31 0.06 1.98 3.07 3.45 1.93 0.48 0.17 100.17
T27 g 65.88 15.43 5.31 0.09 2.25 3.19 3.12 2.21 0.61 0.31 99.65
Loc4 g 66.32 15.28 5.82 0.10 2.28 2.09 2.69 2.38 0.68 0.20 99.85
N7 g 70.94 13.29 3.58 0.10 1.52 3.56 2.89 1.29 0.49 0.14 99.29
N17 g 66.91 15.17 5.26 0.08 1.90 1.08 2.18 2.76 0.68 0.22 99.86
N38c mv 59.53 18.04 8.01 0.23 3.12 4.07 3.35 1.83 0.84 0.20 99.93
N39 mv 57.22 16.77 12.57 0.30 3.46 4.30 1.80 1.81 0.99 0.24 100.24
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a

Rocktypes: p = phyllite, s = schist, g = gneiss, mv = metavolcanic.

4.1. Major phase mineralogy and bulk rock compositions

The samples from the phyllite zone are composed of quartz, plagioclase, biotite, muscovite, chlorite and epidote (Table 1). All samples from the schist zone contain an aluminosilicate (andalusite and/or sillimanite), together with quartz, plagioclase, biotite, muscovite ± garnet (Table 1). The samples of the gneiss zone show variable mineralogy. There are (1) migmatitic gneisses with quartz, biotite, plagioclase, ± sillimanite, ± garnet, ± muscovite, (2) retrogressed migmatitic gneisses with wide-spread late muscovite and chlorite (T15a, N7, N17), and (3) garnet–amphibole-bearing gneisses derived from mafic volcanic rocks (T7, N38, N39; Table 1).

In general, the metasedimentary samples from all three metamorphic zones cover approximately the same range in bulk rock composition (Fig. 2, Table 2). They have compositions with Al2O3, Fe2O3 and MgO similar to average pelites (Fig. 2a; e.g. Shaw, 1956, Spear, 1993) but with generally higher CaO, Na2O and lower K2O values (Table 2). They probably have been derived from pelitic, marly and psammitic protoliths typical for the accretionary prism of the Chugach and Prince William terranes. Sample B40 has a considerably different composition, with very low SiO2 but high Al2O3, Fe2O3, MgO, K2O and TiO2 (Table 2). Samples N7 and B45 have relatively high SiO2 contents. The three garnet–amphibole-bearing metavolcanic gneisses have lower SiO2 and higher Fe2O3 contents than the metasedimentary samples (Table 2).

Fig. 2.

Fig. 2

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(a) AFM and Al2O3–CaO + Na2O–K2O diagrams for bulk rock compositions of the investigated samples. The bulk rock analyses are given in Table 2. (b) CaO vs. Al2O3 content of our samples relative to average metapelite of Shaw (1956). Monazite versus allanite lines according to Wing et al. (2003) for samples with PT conditions similar to the phyllite and schist zone samples. There is no corresponding line for samples with PT conditions similar to the gneiss zone samples.

4.2. Distribution of monazite, allanite and xenotime

Depending on the occurrence of monazite, allanite and ± xenotime, the investigated samples can be classified into four different groups that broadly correspond to the metamorphic zones (Fig. 1 and Table 1).

4.2.1. Phyllite zone: allanite ± xenotime (Fig. 1, Table 1)

All samples from the phyllite zone lack monazite and contain allanite ± xenotime. Allanite occurs either as large single crystals (~ 20–50 μm; Fig. 3a) or as irregular small patches (~ 1 μm) or rims on epidote (Fig. 3b). Xenotime was observed in two of the six phyllite samples (Table 1) as very small crystals (~ 5–10 μm) in the matrix.

Fig. 3.

Fig. 3

Fig. 3

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BSE images of monazite and allanite occurrences within the four different sample groups identified within the CMC. In the phyllite zone, only allanite is present, either as large crystals (a) or small patches (b). In the schist zone, small monazite grains are present in the matrix (c), associated with apatite (d), in garnet or andalusite (e) and associated with allanite (f). In the migmatitic metasedimentary rocks of the gneiss zone, large monazite grains are present in the matrix (g) and are sometimes surrounded by allanite fibres (h). In retrogressed migmatitic gneisses and metavolcanic gneisses, monazite grains are generally small with a relic appearance (i, inset), and overgrown by allanite and epidote (i), or apatite + allanite + epidote coronas (j). Allanite occurs in muscovite–biotite–chlorite shear zones (k, l). Allanite occurs in the matrix of amphibole + garnet-bearing metavolcanic samples (m, n).

4.2.2. Schist zone: small monazite, xenotime ± minor allanite (Fig. 1, Table 1)

In the schist zone samples, monazite crystals are abundant, euhedral and small: the majority are ~ 5–20 μm in size with a few being ~ 20–50 μm in size. Most grains lie in the matrix (Fig. 3c) and some grains are associated with apatite (Fig. 3d). Monazite also occurs as inclusions in andalusite and garnet (Fig. 3e). About 70% of the grains are elongated parallel to the main foliation (regionally termed S2, Pavlis and Sisson, 1995, Scharman et al., 2011) and some are folded by F3 (Fig. 3c). Allanite occurs in four of the nine samples as small fibres/patches surrounding monazite and/or monazite–apatite pairs (Fig. 3f). Xenotime occurs as a minor phase in all samples, i.e. only a few crystals per thin section. The xenotime crystals are small (~ 5–20 μm) and occur mainly in the matrix and sometimes associated with zircon.

4.2.3. Metapelitic gneiss zone: large monazite, xenotime ± minor allanite (Fig. 1, Table 1)

Eight samples from the gneiss zone belong to this group. They are metapelitic gneisses that lack significant retrogression and contain quartz, biotite, plagioclase, ± sillimanite, ± garnet, and ± muscovite (Table 1). Compared to the schist zone samples, the monazite grains in this group are larger (~ 10–100 μm, with a majority in the ~ 20–50 μm range) and have a sub-euhedral crystal shape. Monazite grains are located at grain boundaries or within biotite. Approximately 10% are associated with apatite (Fig. 3g). Approximately 25% of the monazite grains in three samples (B23, KB2, T27) have irregular grain boundaries and are surrounded by fine-grained allanite fibres intergrown with biotite (Fig. 3h). Xenotime is present in all samples and occurs as small crystals (~ 5–15 μm) in the matrix.

4.2.4. Heterogeneous gneiss zone: allanite ± xenotime ± minor monazite (Fig. 1, Table 1)

The remaining eight samples from the gneiss zone belong to this group. The group is mineralogically heterogeneous and includes a qtz + bt + pl gneiss (B21), four retrogressed gneisses with late muscovite ± chlorite (garnet-amphibole bearing sample: T7; metapelitic samples: T15, N7, N17), and two amphibole–garnet-bearing gneisses (N38, N39, Table 1). All samples contain relatively large allanite grains which locally overgrew relict monazite. In the five samples with late muscovite ± chlorite, monazite is present in the matrix as resorbed small crystals (< 50 μm) with irregular grain boundaries, suggesting that these crystals represent relict grains (Fig. 3i, inset). A few monazite grains are overgrown by allanite (Fig. 3i) or complex allanite–apatite–epidote coronas (Fig. 3j). Single, relatively large allanite crystals (100–200 μm) are present within muscovite–chlorite-bearing zones (Fig. 3k), indicating retrograde growth of allanite together with chlorite and muscovite. In sample B21, the small metastable monazite crystals are overgrown by allanite (Fig. 3l), but further indications for retrograde overprint such as muscovite and chlorite are missing in this sample. In the two amphibole-garnet bearing samples, monazite was only observed as two small (< 20 μm), relict crystals in N38. Allanite is wide-spread in both samples as large and strongly zoned crystals in the matrix (~ 20–300 μm in size; Fig. 3m, 3n), which are commonly associated with epidote and plagioclase.

4.3. Chemical zoning and composition of monazites

Because monazite is particularly stable and common in samples from the schist zone and unretrogressed migmatites from the gneiss zone, we further investigated the chemical zoning and composition of these monazites. We conducted high-resolution BSE imaging, element mapping and quantitative chemical analyses on monazites from representative samples of these two groups.

4.3.1. Schist zone

High-contrast BSE images of all small monazite grains revealed that depending on the sample, 60–100% of the grains in each sample show no zoning, whereas the remaining 0–40% of the grains show two or three weak zones of BSE-brightness, corresponding to mass contrasts. Grains from sample B39 that showed most variations in BSE-brightness were selected for element mapping and chemical analysis. Element X-ray maps for Y and Th show only weak variations (Fig. 4a). Representative chemical analyses from different zones of these grains are given in Table 3 and plotted in Fig. 5.

Fig. 4.

Fig. 4

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Representative BSE images and Y and Th element maps of monazites from (a) sample B39 (schist zone, small, unzoned monazites), and (b–c) sample T27 (gneiss zone, large, complexly zoned monazites). Representative monazite EMP analyses (numbered points) are given in Table 3, Table 4.

Table 3.

Representative monazite EMP analyses from sample B39.

B39_1_1 B39_1_2 B39_1_3 B39_2_1 B39_2_2 B39_3_1 B39_3_2 B39_3_3 B39_8_1 B39_8_2 Mina Maxa Ava
SiO2 0.27 0.24 0.42 0.33 0.23 0.30 0.31 0.58 0.29 0.36 0.21 0.81 0.34
Y2O3 1.86 2.28 1.60 1.69 2.50 2.08 1.60 1.50 1.78 1.65 0.71 2.53 1.80
P2O5 29.92 30.03 29.95 29.65 30.04 29.63 29.64 29.44 30.02 29.62 28.60 30.10 29.70
Ce2O3 29.51 28.64 28.38 27.85 28.08 28.56 28.62 29.45 28.86 28.58 27.81 34.70 28.99
ThO2 3.86 4.39 5.87 5.78 5.43 3.84 4.79 3.56 4.30 4.88 2.18 6.27 4.45
Dy2O3 0.21 0.34 0.26 0.57 0.54 – 0.22 0.03 0.33 0.29 – 0.59 0.33
La2O3 13.29 12.71 12.74 12.05 12.64 12.67 13.04 13.48 12.77 12.99 12.05 16.12 13.13
CaO 0.78 0.92 1.10 1.04 1.10 0.84 0.89 0.70 0.85 0.89 0.31 1.14 0.87
Sm2O3 2.12 2.18 2.10 2.46 2.08 2.08 2.14 2.19 2.21 2.33 0.98 2.46 2.12
Pr2O3 3.52 3.26 3.35 3.26 3.15 3.39 3.35 3.42 3.36 3.51 3.15 3.95 3.41
Gd2O3 1.77 1.67 1.93 2.05 1.67 1.73 1.80 1.79 1.88 1.94 0.54 2.20 1.72
Nd2O3 11.83 11.44 11.42 11.99 11.56 11.39 11.35 11.65 12.08 11.72 10.65 12.22 11.55
Eu2O3 0.74 0.72 0.76 0.84 0.76 0.61 0.82 0.74 0.71 0.78 0.61 1.15 0.77
Ho2O3 0.50 0.45 0.44 0.53 0.51 0.51 0.45 0.57 0.56 0.57 – 0.83 0.48
Tb2O3 – 0.02 – 0.01 0.05 0.06 0.12 0.04 0.08 0.02 – 0.17 0.05
Er2O3 – 0.02 – – 0.08 – – – – – – 0.12 0.02
Total 100.16 99.31 100.31 100.10 100.44 97.69 99.14 99.14 100.08 100.14 97.69 101.40 99.72
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a

Minimum, maximum and average compositions from 38 measurements on B39. All analyes are displayed in Fig. 5.

Fig. 5.

Fig. 5

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Compositional diagrams for monazite compositions analysed in samples B39 and T27. Representative EMP analyses are given in Table 3, Table 4.

4.3.2. Gneiss zone

The large monazite grains present in the non-retrogressed migmatitic metasedimentary gneisses have more pronounced zoning than the grains from the schist zone. Up to 80% of the grains within each sample show two or more different BSE zones. Sample T27 showed well-developed zoning and was therefore selected for element X-ray mapping (Fig. 4) and chemical analysis (Fig. 5 and Table 4). The maps revealed that the strong BSE zoning is the result of inverse variations in Th and Y content, e.g. bright BSE zones correspond to high Th and low Y content. The compositional analyses show that the bright BSE zones have high but variable Th, and high La, Ce and Pr contents, whereas the dark zones have high Y, Sm, Nd, and Gd and low Th, La and Ce contents (Table 4, Fig. 5). Bright BSE zones commonly define cores and the dark zones commonly define rims (Fig. 4b). However the zoning is in many cases very patchy or mottled and no clear core-rim structure is present (Fig. 4c).

Table 4.

Representative monazite EMP analyses from sample T27.

T27_1_1 T27_1_2 T27_2_1 T27_2_2 T27_3_1 T27_3_2 T27_3_3 T27_4_1 T27_4_2 T27_4_3 Mina Maxa Ava
SiO2 0.63 0.31 0.69 0.44 0.69 0.40 0.38 0.75 0.32 0.35 0.27 0.93 0.49
Y2O3 0.20 2.12 0.26 2.19 0.44 2.15 2.52 0.60 1.66 3.00 0.06 3.02 1.38
P2O5 29.15 28.70 28.95 29.61 29.06 29.31 29.49 29.34 29.93 30.24 27.77 30.24 29.43
Ce2O3 30.37 28.17 29.56 27.85 30.58 28.02 28.01 28.82 29.30 27.01 27.01 31.56 28.86
ThO2 6.25 4.00 7.55 4.46 5.26 4.21 3.70 7.10 4.05 3.94 1.97 8.76 5.28
Dy2O3 0.00 0.28 0.00 0.00 0.19 0.43 0.29 0.30 0.29 0.50 0.00 0.55 0.24
La2O3 14.31 11.66 13.96 12.13 14.43 12.10 12.09 14.11 12.34 11.80 11.57 15.15 13.11
CaO 1.00 1.05 1.14 1.09 0.90 1.07 1.02 1.17 0.94 1.10 0.83 1.41 1.08
Sm2O3 1.68 2.14 1.72 1.93 1.81 2.15 2.16 1.72 2.23 2.31 1.50 2.60 1.95
Pr2O3 3.63 3.18 3.59 3.04 3.46 3.21 3.16 3.64 3.40 3.12 3.04 3.81 3.40
Gd2O3 0.98 1.58 0.92 1.55 1.14 1.44 1.68 1.18 1.41 1.80 0.73 1.97 1.32
Nd2O3 11.05 11.93 11.58 11.87 11.87 12.58 12.77 10.94 13.02 12.32 10.53 13.55 11.92
Eu2O3 0.08 0.08 0.09 0.08 0.08 0.08 0.08 0.09 0.10 0.09 – – –
Ho2O3 0.17 0.58 0.21 0.37 0.29 0.50 0.59 0.20 0.40 0.60 – 0.68 0.32
Tb2O3 – 0.01 – 0.03 – 0.08 – 0.11 – 0.03 – 0.11 0.02
Er2O3 – 0.02 – 0.08 – 0.11 0.11 – – 0.03 – 0.11 0.02
Total 99.50 95.83 100.21 96.71 100.20 97.84 98.02 100.07 99.39 98.23 95.17 100.95 98.90
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Minimum, maximum and average compositions from 35 measurements on T27. All measurements are plotted in Fig. 5.

4.4. SHRIMP U–Pb dating of monazites from the gneiss zone

Four samples from the gneiss zone that contained larger monazite grains (B21, KB5, KB1 and T27, Table 1) were selected for U–Pb SHRIMP dating. Unfortunately, an attempt of dating the small monazites from three schist samples (B39, Tk9, Tk10) was unsuccessful because mineral separation was unable to recover such small grains, and their low abundance prevented obtaining representative dates by dating them in thin section.

The four samples from the gneiss zone all yielded > 50 grains of monazites per sample, typically ~ 50–200 μm in size. BSE images of separated grains (Fig. 6) confirm the complex zoning and the tendency to bright BSE cores and dark BSE rims that was observed in the thin section of sample T27 (Fig. 4). We performed 13–18 analyses over 10–11 crystals from each sample covering zones with different BSE brightness (core and rims, respectively, Fig. 6, Fig. 7). In sample KB1, one discordant analysis was excluded from the age calculation (discordant analyses are only shown in Table 5, not in Fig. 7). In sample T27, two analyses are excluded due to large errors on the isotopic ratio (Table 5). The remaining analyses from all samples form tight clusters, allowing the calculation of weighted average dates (Fig. 7). The dates vary between 51.0 ± 1.1 Ma and 53.4 ± 1.1 Ma (Fig. 7). The dates of monazite in samples KB1, KB5 and T27 are within error of each other, whereas the mean date of monazite rims from sample B21 (51.0 ± 1.1) is slightly younger than the mean dates of monazite core and rim analyses of sample KB5, but still within error of the KB1 and T27 dates (Fig. 7). Notably, within each sample, core and rims dates are within error of each other.

Fig. 6.

Fig. 6

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Representative BSE images of dated monazite grains from the four gneiss samples. The location of the SHRIMP analyses is indicated by the ellipses and the corresponding 206Pb/238U dates (in Ma ± 1σ) are reported. The scale is the same for all grains. Note that domains with different BSE brightness do not show differences in age. Full U–Pb analyses are given in Table 5.

Fig. 7.

Fig. 7

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(a–d) Concordia diagrams for monazites from the four gneiss samples analysed by SHRIMP. White ellipses correspond to core analyses, grey ellipses correspond to rim analyses. The values in the boxes are weighted average 206Pb/238U ages with corresponding MSWD (mean square weighted deviation). Errors are at the 95% confidence interval. N denotes number of analyses included in age calculation. Only analyses included in the age calculation are plotted (see Table 5 for the complete list of analyses). Core analyses can be correlated with monazite generation Mnz2 and rim analyses with monazite generation Mnz3 based on BSE pattern. (e) Weighted mean ages of rim and core analyses from all samples. Ages overlap within error with the only exception of monazite rims from sample B21, which are slightly younger than the ages of sample KB5.

5. Discussion: factors controlling monazite stability within the CMC

Based on the data on monazite distribution, chemical composition and U–Pb dates presented above, we propose and discuss a model for the behaviour of monazite in the greenschist facies phyllites to anatectic gneisses of the Chugach Metamorphic Complex (Fig. 8). We interpret our observations as the result of three key stages in the evolution of monazite: the monazite-in reaction (Mnz1), the dissolution–precipitation of monazite during partial melting and subsequent cooling (Mnz2 and Mnz3), and finally the breakdown of monazite during retrogression (Fig. 8).

Fig. 8.

Fig. 8

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Schematic diagram displaying the different stages of monazite growth and dissolution observed in the metamorphic zones of the CMC relative to metamorphic grade. In the phyllite zone, only detrital and metamorphic allanite is present. At ~ 550 °C and < 3.4 kbar, monazite growth (Mnz1, medium Y-content) is widespread in the schist zone samples, prior to and contemporaneous with garnet and andalusite growth and parallel to the S2 foliation. In the gneiss zone, Mnz2 growth (Th-rich, Y-poor) occurred due to recrystallisation from partial melts, and was overgrown by Mnz3 rims (Y-rich) during initial cooling and resorption of garnet. In samples located in or close to late muscovite- and chlorite-bearing shear zones, monazite was dissolved and overgrown by allanite and allanite + epidote + apatite coronas.

5.1. Monazite-in reaction at ~ 550 °C and ≤ 3.4 kbar

Within the greenschist facies phyllites, detrital (large, broken crystals, Fig. 3a) and probably metamorphic (small, patchy crystals, Fig. 3b) allanite is present (Fig. 8). At this low grade, monazite was not observed. Therefore, the small monazites present in the schist zone (Fig. 3c–f) are interpreted as entirely metamorphic, implying that a monazite-in reaction occurred at the transition from the phyllite to the schist zone (Mnz1 growth, Fig. 8). The highest-grade sample of the phyllite zone is sample B45, which recorded a peak temperature of 547 ± 50 °C (Table 1, Bruand, 2011) at pressures of ~ 2–3 kbar (Sisson et al., 1989). The lowest-grade sample of the schist zone in which monazite was detected is sample Tk12, which recorded a maximum temperature of 543 ± 23 °C at 3.4 ± 0.9 kbar (Table 1, Bruand, 2011). The bulk rock composition of both samples is very similar, suggesting that the two samples are comparable (Table 2). Whereas sample B45 only contains quartz, plagioclase, muscovite and biotite as major phases, andalusite and garnet occur in addition in sample Tk12. These observations suggest that the monazite-in reaction occurred in these rocks at ~ 550 °C close to or at the andalusite- and garnet-in reactions at a pressure of ≤ 3.4 kbar (Fig. 8).

A monazite-in reaction associated with the andalusite- and garnet-in isograds in metamorphosed pelitic rocks has been reported by numerous workers (e.g. Janots et al., 2008, Kim et al., 2009, Kingsbury et al., 1993, Kohn and Malloy, 2004, Smith and Barreiro, 1990, Tomkins and Pattison, 2007, Wing et al., 2003) and has recently been modelled by Spear (2010). However, the exact reactions leading to monazite growth and the role of the major silicate phases versus other accessory phases during these reactions are debated. Most studies agree that the REE necessary for formation of the monazite probably come from the breakdown of allanite (e.g. Janots et al., 2008, Smith and Barreiro, 1990, Tomkins and Pattison, 2007, Wing et al., 2003). Kohn and Malloy (2004) additionally suggested that the major silicate phases (especially micas and garnet) could liberate considerable amounts of REE to produce monazite. However, Corrie and Kohn (2008) concluded that REEs liberated by major silicate phases are not sufficient and a REE-rich precursor such as allanite is required for monazite formation. In our case, only mica and no garnet is present in the phyllites, indicating that the REE necessary for monazite formation were probably derived from a REE-rich precursor such as allanite and/or xenotime. However, allanite to monazite break-down textures such as described by Janots et al. (2008), Kim et al. (2009), Tomkins and Pattison (2007) or Wing et al. (2003) were not observed in our samples. This may be due to the high level of equilibration and poor preservation of prograde relics in this sequence.

Aluminosilicates have been proposed to be necessary for the formation of monazite at or close to the andalusite and garnet-isograds. Wing et al. (2003) proposed the following reaction for the formation of monazite at the aluminosilicate-in isograd: 3 allanite + 1 apatite + 1 muscovite + 4 aluminosilicate + 3 quartz = 3 monazite + 1 biotite + 8 plagioclase + 2 H2O. However, Tomkins and Pattison (2007) observed that the monazite-in reaction in their rocks occurred slightly down temperature of the aluminosilicate-in reaction and proposed the following general formation mechanism: allanite + apatite = monazite + plagioclase + magnetite. Similarly, Janots et al. (2008) did not observe a clear link with the aluminosilicate-in reaction and proposed the following generalised prograde reaction: allanite + apatite + Al–Fe–Mg phases1 = monazite + anorthite + Al–Fe–Mg phases2 (e.g. biotite, muscovite, garnet). In the Chugach Metamorphic Complex, the rare observation of monazite inclusions in andalusite and garnet (Fig. 3e) indicates that the monazite-in reaction probably occurred before the formation of andalusite and garnet, similar to observations by Tomkins and Pattison (2007). In addition, monazite grains in sample B39 from the schist zone are relatively Y-rich (1–2.5 wt.% Y2O3, Fig. 5), indicating that they probably grew before garnet, which preferentially incorporates Y (e.g. Foster and Parrish, 2003). Based on these qualitative observations, we propose that probably a reaction similar to the one proposed by Janots et al. (2008) or Tomkins and Pattison (2007) was responsible for monazite formation in the schist zone samples.

Various workers have shown that the Ca-content of the bulk rock and/or of the fluid phase influence the stability of monazite (e.g. Budzyn et al., 2011, Foster and Parrish, 2003, Janots et al., 2008, Spear, 2010, Wing et al., 2003). Several indicators have been used to show this influence. Wing et al. (2003) used CaO versus Al2O3 plots of andalusite-, kyanite- and sillimanite-zones to show that allanite is stable instead of monazite in their samples if the bulk rock contains 2–3 times the average pelite CaO-content. For the CMC samples of similar metamorphic grade (the phyllite and schist zones, Fig. 2b) CaO versus Al2O3 plots do not display such a relationship (Fig. 2b). For example, samples B39 and B40, which both lie in the stability field for allanite as defined by Wing et al. (2003), contain abundant stable monazite, and allanite is only present as retrograde fibres or as relict prograde crystals. These retrograde allanite fibres are not restricted to samples from the allanite stability field of Wing et al. (2003), but also occur in samples in the monazite stability field (Fig. 2). Foster and Parrish (2003) proposed an equation that links the temperature of monazite growth to the CaO/SiO2-ratio of the rock. This equation gives monazite-in temperatures of ~ 670 °C and ~ 760 °C for samples B39 and B40, respectively, suggesting that allanite should instead be stable in these samples, which is not the case. Therefore, for the CMC samples, CaO versus Al2O3 plots or the CaO/SiO2 ratio seem not to be the tools for revealing the Ca-influence on monazite stability. Janots et al. (2008) suggested that the CaO/Na2O ratio of a sample might influence the breakdown of allanite to monazite between temperatures of ~ 500–650 °C, with allanite being stable to higher temperatures with higher CaO/Na2O ratio. The CaO/Na2O ratio of all our schist samples is below 1.1, which, according to Janots et al. (2008), is clearly in the field where allanite breaks down to monazite at temperatures of ~ 550 °C–600 °C. This is in agreement with stable monazite being present in all schist zone samples. However, since the CaO/Na2O ratio in our samples does not vary greatly, we could not investigate whether this ratio correctly describes monazite vs allanite stabilities in our rocks.

5.2. Dissolution and precipitation of monazite during partial melting and initial cooling

The monazite-bearing samples from the gneiss zone contain abundant, large monazite grains that are much more zoned in BSE images than the monazite grains from the schist zone. Element X-ray mapping and chemical analyses from one representative sample (T27) revealed the presence of two compositional domains: Th- and LREE-rich and Y-poor domains versus Y-, MREE-rich, and Th- and LREE-poor domains (Fig. 5). The distribution of the domains is mainly patchy, with a tendency of the Th-rich, Y-poor domains constituting cores and the Th-poor, Y-rich domains constituting rims (Fig. 4, Fig. 6). No differences in U–Pb dates were detected between the different zones (Fig. 6, Fig. 7). Complex compositional zoning of metamorphic monazite in upper amphibolite-facies rocks and in migmatites has been reported by many workers (e.g. Berger et al., 2009, Fitzsimons et al., 1997, Foster et al., 2002, Foster et al., 2004, Gibson et al., 2004, Kelsey et al., 2008, Kohn et al., 2005, Pyle and Spear, 2003, Pyle et al., 2001, Rubatto et al., 2001). Zoning in Y has been investigated extensively because Y is also incorporated in other phases such as garnet and xenotime (e.g. Pyle et al., 2001, Zhu and O'Nions, 1999a, Zhu and O'Nions, 1999b). The partitioning of Y between the different minerals may permit conclusions regarding the timing of monazite growth relative to garnet and/or xenotime growth (e.g. Pyle et al., 2001). Monazite grains in the schist zone (B39) are relatively Y-rich (Fig. 5), possibly indicating growth prior to garnet formation. Monazite grains in the gneiss zone (T27), however, show Y-poor zones preferentially in the grain cores (Fig. 4). In addition, these cores show complex zoning in Th (Fig. 4). The composition of these cores suggests that they grew when garnet was already present in the assemblage (e.g. Foster and Parrish, 2003). They could therefore represent a younger generation of monazite than the Y-rich monazite in the schist zone (Mzn2 growth, Fig. 8). This would imply that Y-rich monazite (Mnz1), which grew before or simultaneously with garnet (i.e., monazite occurs as inclusions in garnet in the schist samples), was dissolved prior to the formation of the Y-poor monazite (Mnz2). The samples containing Mnz2 experienced partial melting under water-saturated conditions (Bruand, 2011) with garnet being stable both before and during the melting stage. Monazite solubility in metapelitic melts under water-saturated conditions can be relatively high (e.g. Montel, 1993, Rapp and Watson, 1986, Zeng et al., 2005) suggesting that dissolution of lower-grade monazite during partial melting could have occurred in the CMC. The Th-rich, Y-poor cores could then have crystallised during melt recrystallisation, with the high Th-content being the result of preferential Th incorporation into monazite in a melt-buffered system (Mnz2 growth, Fig. 8; e.g. Kohn and Malloy, 2004, Kohn et al., 2005). The recrystallisation of these large monazite grains under peak metamorphic conditions (~ 650–700 °C and ~ 4–9 kbar; Bruand, 2011) is in agreement with experimental monazite stability fields in rocks of similar bulk composition (e.g. Janots et al., 2007). Similar monazite zoning has also been described during water-absent biotite-breakdown melting (e.g. Dumond et al., 2010).

Y-rich zones/rims similar to the Y-rich zones of sample T27 (Fig. 4) have been described by other workers from migmatitic rocks, and have generally been interpreted as representing a monazite generation which grew during breakdown of garnet and/or in equilibrium with xenotime (e.g. Foster and Parrish, 2003, Kohn et al., 2005, Mahan et al., 2006, Pyle and Spear, 2003). Sample T27 contains garnet with irregular grain boundaries that show a strong outward increase in Mn and a slight decrease in Mg in the rim (Bruand, 2011). Sharp rimward zoning in Mg, Fe or Mn in otherwise homogenous garnet is commonly attributed to diffusional exchange and/or net transfer reactions between garnet and biotite during cooling (e.g. Kohn et al., 2001, Spear, 1993). Irregular grain boundaries and truncated zoning together with Mn-rich rims are commonly interpreted as resorption of garnet during which liberation of Y may occur. Additionally, dissolution of xenotime, which is indeed rare in these samples, in the final stage of melting would liberate Y. Therefore we suggest that the high-Y monazite rim (Mnz3 growth, Fig. 8) occurred after release of Y during garnet resorption and/or xenotime dissolution (e.g. Foster and Parrish, 2003, Kohn et al., 2005, Spear, 1993).

5.3. Break-down of monazite in retrograde fault zones

The retrogressed gneisses contain only small, relict monazite grains and large allanite crystals. In samples T7, T15a, N7 and N17, these allanite crystals are texturally linked to the occurrence of late muscovite and/or chlorite, which is commonly concentrated along foliation-parallel zones (Fig. 3k, Fig. 8). These muscovite–chlorite bearing zones are wide-spread in the southern and south-eastern parts of the CMC, and contain a down-dip lineation with N-side up sense of movement (Gasser et al., 2011). The zones occur near and parallel to major fault zones which bound the complex: the Contact fault system in the west and the Fairweather/Art Lewis Fault System in the south-east (Fig. 1; Gasser et al., 2011, Pavlis and Roeske, 2007, Plafker et al., 1994). These fault systems were responsible for exhumation of the CMC during the Eocene to Neogene (Gasser et al., 2011). We therefore interpret the allanite growth in these samples as a consequence of fluid-assisted greenschist facies retrogression associated with exhumation of the gneisses along the bounding fault zones (Fig. 8). Similar allanite–apatite corona growth around monazite (as observed in samples N7 and N17, Fig. 3j) has been described by Finger et al. (1998) as a result of an amphibolite-facies overprint of granitoids and by Krenn et al. (2008) as a result of a greenschist-facies overprint of amphibolite-facies metapelites. Budzyn et al. (2011) produced similar coronas in experiments at 4.5–6.1 kbar, 450–500 °C and with high Ca-fluids, suggesting that the fluid composition plays a major role in development of such corona growth.

The two amphibole–garnet-bearing metavolcanic gneisses N38 and N39 appear to represent exceptions: the large allanite crystals present in these samples do not appear to be texturally related to the late muscovite–chlorite zones, and we infer that they represent part of the peak metamorphic assemblage. Allanite, plagioclase and garnet in these samples show complex zoning patterns in Ca, suggesting a complicated reaction history for this assemblage within these samples. The high molar Ca-content of these samples (~ 5 Ca mol%) relative to the other investigated samples (~ 1–3.8 Ca mol%), probably in combination with slightly higher overall peak pressures in the south-eastern part of the CMC (e.g. pressure controls the monazite-vs. epidote/allanite-stability field as depicted in Fig. 6 in Janots et al., 2007), likely favoured stability of allanite at the metamorphic peak in these samples.

5.4. Remarks on the U–Pb ages and regional implications

U–Pb SHRIMP dates from core and rim domains can be correlated with monazite generations Mnz2 and Mnz3, respectively, based on their similar BSE pattern (Fig. 6). Single spot analyses from rim and core domains from the different samples range from 50.3 ± 1.9 Ma to 54.8 ± 0.9 Ma (Table 5), whereas the weighted mean dates of rim and core domains of the four samples range from 51.0 ± 1.1 Ma to 53.4 ± 1.1 Ma (Fig. 7). There is no systematic difference between analyses from core and rim domains in the different samples (Fig. 7e). This indicates that Mnz2 and Mnz3 growth occurred during a short period of time of less than 5 m.y., below the resolution of our dating method (Fig. 8). For samples KB5 and T27, U–Pb SHRIMP ages of metamorphic zircon rims are within error of the monazite ages for the same samples (54.0 ± 0.8 Ma for KB5 and 52.6 ± 0.8 Ma for T27, Gasser et al., in press). Four additional U–Pb SHRIMP ages of metamorphic zircon rims from the CMC and one age from a granodioritic intrusion also occur within error of the monazite ages (Gasser et al., in press). Metamorphic zircon rims were interpreted as being related to the partial melting event, and we suggest that monazite formation probably occurred during the same time period as the zircon-rim formation close to or at the metamorphic peak. Together, zircon and monazite ages indicate a short time period from ~ 55 to ~ 50 Ma during which partial melting occurred in the CMC, contemporaneously for at least ~ 300 km along strike (Gasser et al., in press).

6. Conclusions

In the Chugach Metamorphic Complex, which ranges from greenschist facies phyllites to upper amphibolites facies migmatites, monazite is affected by three major events: (1) At ca. 550 °C and ≤ 3.4 kbar, the detrital and metamorphic allanite present in the greenschist facies phyllites broke down to form medium-Y monazite (Mnz1). The widespread growth of monazite occurred prior to and simultaneously with garnet and andalusite growth, probably following the general reaction allanite + apatite + Al–Fe–Mg phases1 = monazite + anorthite + Al–Fe–Mg phases2. Aluminosilicates or garnet probably did not take part in the monazite-forming reaction, similarly to what has been described by Janots et al. (2008) and Tomkins and Pattison (2007). CaO versus Al2O3 plots, CaO/SiO2 and CaO/Na2O ratios do not reveal the influence of Ca on the monazite-in reaction in our samples. (2) During partial melting of the gneisses, the early medium-Y monazite generation Mnz1 was apparently dissolved, and a second, Y-poor, Th-rich monazite generation Mnz2 crystallised, which we interpret to have formed synchronous with or shortly after the metamorphic peak of ~ 650–700 °C and 4–9 kbar. During initial cooling and resorption of garnet rims, a third, Y-rich, Th-poor monazite generation Mnz3 formed. U–Pb SHRIMP dating of the second and third monazite generations in the gneisses shows that Mnz2 and Mnz3 are indistinguishable in age, and that the two monazite generations crystallised during a short time span of ≤ 5 m.y. at ~ 55–50 Ma. (3) Retrogression associated with the development of muscovite- and/or chlorite-bearing shear zones destabilised monazite and led to the formation of allanite and/or allanite + epidote + apatite coronas. The combination of microstructural observations, petrology, trace element geochemistry and geochronology allowed different monazite generations to be related to different metamorphic stages. The results illustrate that polyphase, complex monazite growth and dissolution can occur during a single, relatively short-lived metamorphic cycle.

Acknowledgements

Financial support by the FWF Austrian Science Fund project P-19366 is greatly acknowledged. J. Neubauer (University of Graz) is thanked for help with the SEM analyses, F. Zaccharini (University of Leoben) is thanked for help with the monazite analyses, U. Klötzli (University of Vienna) is thanked for initially suggesting to study monazite in the CMC. Gregory Dumond and an anonymous reviewer are thanked for very constructive and useful reviews.

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