Subduction and exhumation mechanisms of ultra-high and high-pressure oceanic and continental crust...

Subduction and exhumation mechanisms of ultra-high andhigh-pressure oceanic and continental crust at Makbal(Tianshan, Kazakhstan and Kyrgyzstan)

M. MEYER,1 R. KLEMD,1 E. HEGNER2 AND D. KONOPELKO3

1GeoZentrum Nordbayern, Universit€at Erlangen-N€urnberg, Schlossgarten 5a, D-91054, Erlangen, Germany([email protected])2Department f€ur Geo- und Umweltwissenschaften and GeoBio-Center, Universit€at M€unchen, Theresienstraße 41, D-80333,Munich, Germany3Geological Faculty, St. Petersburg State University, 7/9 University Embankment, St. Petersburg 199034, Russia

ABSTRACT The Makbal Complex in the northern Tianshan of Kazakhstan and Kyrgyzstan consists ofmetasedimentary rocks, which host high-P (HP) mafic blocks and ultra-HP Grt-Cld-Tlc schists(UHP as indicated by coesite relicts in garnet). Whole rock major and trace element signatures ofthe Grt-Cld-Tlc schist suggest a metasomatized protolith from either hydrothermally altered oceaniccrust in a back-arc basin or arc-related volcaniclastics. Peak metamorphic conditions of the Grt-Cld-Tlc schist reached ~580 °C and 2.85 GPa corresponding to a maximum burial depth of ~95 km.A Sm-Nd garnet age of 475 � 4 Ma is interpreted as an average growth age of garnet during pro-grade-to-peak metamorphism; the low initial eΝd value of �11 indicates a protolith with an ancientcrustal component. The petrological evidence for deep subduction of oceanic crust poses questionswith respect to an effective exhumation mechanism. Field relationships and the metamorphic evolu-tion of other HP mafic oceanic rocks embedded in continentally derived metasedimentary rocks atthe central Makbal Complex suggest that fragments of oceanic crust and clastic sedimentary rockswere exhumed from different depths in a subduction channel during ongoing subduction and arenow exposed as a tectonic m�elange. Furthermore, channel flow cannot only explain a tectonicm�elange consisting of various rock types with different subduction histories as present at the centralMakbal Complex, but also the presence of a structural ‘dome’ with UHP rocks in the core (centralMakbal) surrounded by lower pressure nappes (including mafic dykes in continental crust) and volu-minous metasedimentary rocks, mainly derived from the accretionary wedge.

Key words: exhumation; m�elange; subduction; Tianshan; (ultra)high-P.

INTRODUCTION

The Tianshan orogenic belt is a Palaeozoic crystallineunit that is part of the Central Asian Orogenic Belt(CAOB), one of the largest, most complex andlong-lasting orogens in the world, which was formed bylateral accretion of magmatic arcs and microcontinents,as well as vertical accretion of mantle-derived mate-rial to the crust (e.g. Jahn et al., 2000a,b; Chen &Jahn, 2004; Xiao et al., 2004; Kr€oner et al., 2007,2008; Windley et al., 2007; Long et al., 2011). TheTianshan extends over more than 2500 km fromUzbekistan via Tajikistan, Kyrgyzstan and Kazakh-stan to northwestern China (e.g. Zonenshain et al.,1990; S�eng€or et al., 1993). The belt is subdivided intothe Southern, Middle and Northern units, and isbounded to the south by the Tarim craton (Fig. 1a).Enclosed within the orogenic belt are a number ofdifferent (U)HP [(ultra)high-P] metamorphic com-

plexes representative of different suture zones, andwhich record different subduction and exhumationmechanisms. Of these, the Makbal Complex (Fig. 1a)is the focus of attention in this paper.Most (U)HP–LT metamorphic complexes of the

Tianshan have been investigated in detail during thelast decades, with the geological, petrological,geochemical and geochronological data providingconstraints on the Palaeozoic tectonothermal evolu-tion of the CAOB (for a review, see Klemd et al.,2014a and references therein). Nevertheless, thegeodynamic settings in which lithotectonic units char-acteristic of the southern CAOB and consisting ofmetasedimentary rocks with interleaved mafic blocksand boudins evolved, and became juxtaposed haveyet to be resolved. The determination of whether thecountry rocks and enclosed incoherent mafic bodieshave a common or different P–T evolution will aidin elucidating the subduction and exhumation

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mechanisms of (U)HP metamorphic rocks responsiblefor these sedimentary-dominated tectonic m�elanges.Additionally, understanding the evolution of the sedi-mentary-dominated m�elange of the Makbal Complexwill help resolve whether this complex has a commonor different origin from the other (U)HP complexesof the Northern Tianshan.

At Makbal in the Northern Tianshan of Kyrgyz-stan and Kazakhstan schistose rocks with mainlyporphyroblastic garnet, chloritoid and talc (in thefollowing termed Grt-Cld-Tlc schists) are often asso-ciated with eclogitic boudins hosted by metasedimen-tary rocks (quartz-mica schists and mica-bearingquartzites). Some Grt-Cld-Tlc schists provideevidence of UHP metamorphism as can be inferredfrom rare coesite inclusions in garnet (e.g. Togonba-eva et al., 2009; Tagiri et al., 2010; Konopelkoet al., 2012; this study). Thus, this Grt-Cld-Tlcassemblage, indicative of UHP, has the potentialfrom petrological and thermodynamic modelling toconstrain the metamorphic conditions and so aid inthe reconstruction of subduction and exhumation

mechanisms accompanying deep subduction. Nocomprehensive analysis of this geodynamicallyimportant Makbal Complex rock is available, andso is used here to constrain protolith composition,metasomatic reactions, the timing of metamorphismand its overall metamorphic evolution.The results of this study were combined with previ-

ously published chemical data for whole rocks, min-eral compositions and P–T reconstruction of othereclogite facies mafic boudins (Meyer et al., 2013),and an exhumation model involving exhumationduring ongoing oceanic subduction and subsequentexhumation involving early continental subduction isproposed.

GEOLOGICAL OVERVIEW

Geological setting of the Makbal Complex within theTianshan

The Makbal Metamorphic Complex is located in theKyrgyz Northern Tianshan and comprises areas on

(a)

(b)

Fig. 1. (a) Tectonic map of the westernTianshan (compiled after Burtman, 2006;Hegner et al., 2010; Glorie et al., 2011 andreferences therein). (U)HP complexes arelabelled Makbal (M), Aktyuz (Ay), Anrakhai(An), Atbashy (Ab) and Akeyazhi (Ak) innorthwestern China. Major fault zones orsutures are the Nikolaev Line or NorthNalati Fault (1), the South Tianshan SutureZone (2), the Djalair-Naiman Suture Zone(3) and the Talas Fergana Fault (4). (b)Schematic geological map of the MakbalComplex (based on Tagiri & Bakirov, 1990;Togonbaeva et al., 2009; modified afterRojas-Agramonte et al., 2013): red circlesindicate sample location of rocksinvestigated in this study; white circles aresample location of previous studies.

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862 M. MEYER ET AL .

both sides of the Kazakh-Kyrgyz border. The wes-tern part of the Tianshan in Kyrgyzstan and Kazakh-stan is subdivided by (south)west–(north)east-strikingmajor tectonic discontinuities (Fig. 1a): The NorthernTianshan is separated from the Middle Tianshan bythe Nikolaev Line (Main Tectonic Line), a LateCarboniferous to Early Permian (Hercynian) strike–slip fault following a Silurian to Early Devonian(Caledonian) suture (e.g. Bakirov et al., 2003; Gaoet al., 2009; Qian et al., 2009; Orozbaev et al., 2010;Konopelko et al., 2012).

The Kyrgyz Northern Tianshan consists of severalPrecambrian continental fragments (interpreted asmicrocontinents), ophiolite belts representing thesutures of former oceanic basins (e.g. Djalair-Nai-man, Kyrgyz-Terskey) as well as HP to UHP eclogitefacies rocks and large volumes of Early Palaeozoicgranitoids all amalgamated prior to the Early Ordo-vician (e.g. Windley et al., 2007; Ghes’, 2008;Alexeiev et al., 2011; Kr€oner et al., 2012, 2013). TheKyrgyz-Terskey Ocean (regarded as part of a largerPaleo-Asian Ocean) between the Northern and Mid-dle Tianshan existed from Late Precambrian tothe early Late Ordovician and was subducted to thenorth (present-day coordinates) underneath thesouthern margin of the North Tianshan (Lomizeet al., 1997; Konopelko et al., 2008). The MakbalComplex is situated near the southwestern margin ofthe Northern Tianshan (Fig. 1b); however, the affilia-tion of the microcontinent southwest of the Makbalarea to either the Middle Tianshan (e.g. Windleyet al., 2007; Qian et al., 2009; Biske & Seltmann,2010; Alexeiev et al., 2011; Seltmann et al., 2011) orthe Northern Tianshan (e.g. Gao et al., 2009; Oroz-baev et al., 2010; Konopelko et al., 2012; Kr€oneret al., 2014) is not clear yet (e.g. discussion in Kr€oneret al., 2013).

The Makbal Complex (Fig. 1b) comprises an areaof ~12 9 25 km, but most recent studies focusedon a small area in the central part where eclogitefacies rocks have been reported in Medvedeva(1960, 1961), Bakirov (1978), Kushev & Vinogradov(1978), and where HP and UHP rocks are exposed(e.g. Tagiri et al., 1995, 2010; Konopelko et al.,2012; Meyer et al., 2013). The central Makbal Com-plex was interpreted to represent a tectonic m�elangeformed by a buoyancy-driven ascent of quartzitesand metapelites that incorporated exotic maficblocks metamorphosed at different times and depthsin a subduction zone (Bakirov et al., 1998; Tagiriet al., 2010).

In the central Makbal Complex (Fig. 1b, for aschematic cross-section, refer to Togonbaeva et al.,2009; Tagiri et al., 2010), mafic blocks and boudinspredominantly occur as tectonic lenses (diameterranging between 0.1 and 50 m) enclosed in (U)HPmetasedimentary rocks (quartz-mica schists) or UHPgarnet-chloritoid-talc schists. Mafic rocks comprisemainly strongly retrogressed eclogites sensu lato

(eclogites s.l., cf. Meyer et al., 2013 for furtherdetails), eclogite facies glaucophane-garnet-omphacitebearing rocks (‘glaucophanite’, see Meyer et al.,2013) and the Grt-Cld-Tlc schists investigated inthis study. Furthermore, garnet amphibolites (e.g.Rojas-Agramonte et al., 2013) were found ~13 kmsouth of the central Makbal area. Rare occurrencesof eclogitic cores of mafic lenses were described byseveral authors (Tagiri et al., 1995, 2010; Togonbaevaet al., 2010; Konopelko et al., 2012); however, noneof the here studied mafic samples contained an eclog-itic (sensu stricto) mineral assemblage.

Previous P–T estimates and ages

UHP Grt-Cld-Tlc schist

This rock type experienced UHP conditions, as indi-cated by coesite inclusions or coesite relicts sur-rounded by palisade quartz inclusions in garnet. PeakP–T conditions are estimated at ~560 °C and2.8 GPa (Tagiri et al., 2010). A U-Pb SHRIMP ageof 502 � 10 Ma was reported for metamorphiczircon rims (Konopelko et al., 2012), whereas Tagiriet al. (2010) obtained a K-Ar phengite age of509 � 13 Ma. In addition, Th-Pb CHIME monaziteages range between 481 � 26 Ma and 480 � 56 Ma(Togonbaeva et al., 2009).

Metasedimentary rocks

Both mafic lenses and Grt-Cld-Tlc schists areembedded in metasedimentary rocks and marbles,the former consisting mainly of quartzites andquartz-mica schists. For these rock types, geother-mobarometrical and/or direct geochronologicalstudies have not been undertaken yet. However,detrital magmatic and metamorphic zircon ingarnet-muscovite-quartz schists, interpreted as theimmediate host rock of the eclogites, has U-PbSHRIMP ages ranging from 3300 to 1900 Ma(Degtyarev et al., 2013).

Mafic lenses

Quartz pseudomorphs after coesite in garnet of theeclogitic lenses have been reported by Tagiri et al.(2010), but were not found in any of our mafic sam-ples. Prograde P–T paths were reconstructed forthree eclogite facies mafic lenses using garnet isopleththermobarometry with peak pressure conditions of~560 °C at nearly 2.4 GPa and ~520 °C at 2.15 GPaand ~550 °C at 2.5 GPa respectively (Meyer et al.,2013). The lack of UHP indicators is in accordancewith a peak P–T estimate of 560 °C at 2.0 GPa foran eclogitic sample with a K-Ar paragonite age of482 � 17 Ma (Tagiri et al., 1995). Additional agesfor metamorphic zircon rims from Makbal HP eclog-ites were reported by Konopelko et al. (2012) with


SUBDUCT ION AND EXHUMATION OF (U )HP CRUST AT MAKBAL 863

U-Pb SHRIMP ages between 509 � 7 Ma and498 � 7 Ma. Two HP garnet amphibolites with peakP–T conditions of ~620 °C at 1.4 GPa, sampled~13 km south of the central Makbal locality, yieldeda Lu-Hf garnet isochron age of 470 � 2.5 Ma thatwas interpreted as the age of prograde garnet growth(Rojas-Agramonte et al., 2013).

ANALYTICAL TECHNIQUES

Whole rock geochemical composition

The major elements were analysed on glass discs byX-ray fluorescence (XRF) at the Geodynamics andGeomaterials Research Division, W€urzburg Univer-sity using a MINIPAL 4 XRF spectrometer (PANa-lytical). Accuracy and precision were better than1–2% (1r) for the major oxides. For calibration,international rock standards were used.

The trace elements were analysed on the sameglass discs at the GeoZentrum Nordbayern, Erlan-gen University by laser ablation inductively coupledplasma mass spectroscopy (LA-ICP-MS) using anAgilent 7500i quadrupole mass-spectrometer coupledwith a New Wave Research 266LUV laser ablationsystem. Trace elements were calculated usingGLITTER version 3.0 data reduction for LA-ICP-MS(Macquarie Research Ltd, 2000). The SiO2 contentof each sample was used as an internal standard.The instrument was calibrated with the glass refer-ence material NIST SRM 612 (values of Pearceet al., 1997). The international basalt standard BE-N was analysed to check the accuracy and repro-ducibility of analyses (cf. Jochum et al., 2006). Eachsample was analysed four times and the reporteddata represent average values. The reproducibilityfor the four analyses is on average better than 5%except rare earth elements (REE), Cs, Ta and Pbwith an average relative standard deviation (RSD)<10%. Laser ablation was performed along lines of~1100 lm with a repetition rate of 20 Hz and anablation crater with 50 lm in diameter. Laserenergy was set at 1.1 mJ with an energy density of55 J cm�2 and a scan velocity of 40 lm s�1. Mea-surement was conducted as time-resolved analysisusing 20 s background and 30 s sample countingtimes. Sample contamination by the flux materialand preparation procedure was monitored with themeasurement of ‘blanks’ and is considered negligiblefor the analysed element concentrations. Furtherdetails regarding the mass spectrometer and ablationand standardizing procedures respectively are givenby Schulz et al. (2006).

Mineral analysis and mapping

The major elements and compositional maps wereobtained at the GeoZentrum Nordbayern, ErlangenUniversity on a JEOL JXA-8200 electron microprobe

equipped with five wavelength-dispersive spectrometers(WDS). For mineral analyses, the microprobe wasoperated with an acceleration voltage of 15 kV, aprobe current of 15 nA and a probe diameter of3 lm. For signal calibration, both natural and syn-thetic mineral standards were used. The matrix cor-rection of the raw counts was conducted using aZAF procedure. Major element maps of garnet weremeasured by EMPA (electron microprobe analysis)in WDS (Ca, Fe, Mn, Mg, Al) and EDS (energy-dispersive spectrometer) mode with an accelerationvoltage of 15 kV, focused beam, 100 ms dwelltimeper pixel and a probe current of 200 nA.The trace element composition of garnet was mea-

sured by LA-ICP MS at the GeoZentrum Nordbay-ern, Erlangen University using a UP193 FX NewWave Research Excimer Laser coupled with aAgilent 7500i ICP-MS. Element concentrations werecalculated using GLITTER software (see above) withthe SiO2 concentration from garnet microprobeanalyses as an internal standard. The NIST SRM612 material (Pearce et al., 1997) was used for exter-nal signal calibration. Data reproducibility and accu-racy were determined on NIST SRM 614 and BCR-2, giving a reproducibility and accuracy for most ele-ments ≤10% (RSD). Laser ablation was performedon single spots with a repetition rate of 15 Hz and acrater diameter of 50 lm (in some cases 35 lm).Laser energy was set at 0.59 GW cm�2 with anenergy density of 2.94 J cm�2. Analyses were per-formed as time-resolved analyses using 20 s back-ground and 25 s sample counting times.

Raman spectroscopy

Raman spectroscopy of coesite inclusions in garnetwas conducted at Deutsches GeoForschungsZentrum(GFZ) Potsdam at the Section of Chemistry andPhysics of Earth Materials. A HORIBA Jobin YvonLabRAM HR800 VIS Raman spectrometer was usedwith an excitation wavelength of 473 nm (diode-pumped solid-state laser). The laser was focused onthe thin section by an Olympus BXFM microscopewith a 509 objective. Rayleigh diffusion was elimi-nated by edge filters, the entrance slit was set at100 lm and the signal was dispersed using a1800 l mm�1 grating. Each measurement was con-ducted three times with an acquisition time of 30 s.Prior to measurement, the spectrometer was cali-brated with a silicon standard.

Sm-Nd garnet dating

The Sm and Nd isotopic analyses on garnet and wholerock material of Grt-Cld-Tlc schist sample 10-16 werecarried out at Munich University following the proce-dures of Hegner et al. (2010).Whole rock and matrix material were crushed and

ground to a fine powder. Garnet grains were separated



from the matrix, crushed and sieved. The grain-sizefraction of 250–160 lm was chosen for analysisbecause it showed least intergrowth of minerals. Thegarnet fraction was washed in water, dried in acetoneand purified using a Frantz magnetic separator. Anoptically clean garnet separate was hand-picked undera binocular microscope. The garnet sample was repeat-edly washed in distilled water, dried in acetone andground in a boron carbide mortar. The powder wasleached in 50% HNO3 at ~70 °C overnight andwashed five times in water. It was then split and analiquot as well as the whole rock and matrix powderswere spiked with a 149Sm-150Nd tracer before digestionin HF-HClO4 at 125 °C on a hotplate for 1 week. Thelight rare earth elements (LREEs) were separated on5 ml quartz columns with a resin bed of AG50 9 12,200–400 mesh. Samarium was separated from Ndusing quartz columns filled with 1.9 ml HDEHP-coated teflon powder. Isotopic compositions weremeasured on a TRITON thermal ionization mass spec-trometer in static data collection mode. 143Nd/144Ndwere normalized to 146Nd/144Nd = 0.7219 and theexternal precision for 143Nd/144Nd is estimated at~5 9 10�6 (2SD). The JNdi-1 reference materialyielded 143Nd/144Nd = 0.512103 � 3 (2SD, n = 8) dur-ing the course of this study.

PETROGRAPHY AND MINERAL CHEMISTRY

UHP Grt-Cld-Tlc schist (10-16)

One sample of the Grt-Cld-Tlc schist (10-16, Fig. 1b)was collected at the central Makbal locality ~5 m awayfrom a mafic boudin (‘glaucophanite’ 14R – fig. 2 inMeyer et al., 2013). To the north, the Grt-Cld-Tlcschist is in contact with a quartzite band (fig. 2 inMeyer et al., 2013). The schist is strongly foliated anddips to the north (010/45). In places, this schist showsevidence of dextral shearing as recorded in garnet por-phyroblasts (Fig. 2b).

Porphyroblastic garnet (~35 vol.%, modal amountbased on point counting in thin section) with a max-imum diameter of 1.5 cm (Fig. 2c,d) has an averagecomposition of Alm78Prp19Grs2Sps1 and containssingle to low amounts of non-oriented homoge-neously distributed inclusions of tiny round rutile,elongated chloritoid (mean: XFe2þ = 0.75), small talc(mean: XMg = 0.90), differently sized and shapedquartz and very rare chlorite (XFe2þ = 0.32). Onegarnet grain contains an inclusion of relict coesite asconfirmed by Raman spectroscopy (Fig. 3a–c) sur-rounded by palisade quartz. Other quartz inclusionswithout coesite relicts are mostly polycrystalline(Fig. 2d) and are surrounded by garnet with radialcracks.

Two different zoning patterns were observed in gar-net porphyroblasts: garnet grains (Grt1) with inclu-sions of chloritoid, quartz and/or coesite relicts, raretalc, very rare chlorite and rutile (Fig. 2c,d) show a

higher concentration of MnO in the core (mean corecomposition: Alm82–84Prp12–13Grs2–3Sps1–2) and somegarnet grains exhibit a well-preserved bell-shaped pro-grade zoning pattern for XSps (Fig. 4). The meaninner mantle (m1) composition is Alm80–84Prp13–17-Grs

2Sps0.6–1.0 while in the outer mantle (m2) and rim

region (mean m2 to rim compositions: Alm70–80-

Prp18–29Grs0.7–2.0Sps0.2–0.3), the spessartine-componentis lower than or similar to that observed for the coreconcentrations of the garnet grains (Grt2) withoutchloritoid, quartz, talc and chlorite inclusions(Fig. 2c,d). Even though the spessartine-component isgenerally low compared with Grt1, the Grt2 grainsstill show prograde zoning patterns with an averagecore composition of Alm80Prp18Grs1.6–1.7Sps0.3–0.4 anda rim composition of Alm73–74Prp25–26Grs0.5Sps0.2–0.3.Both garnet types are characterized by a high XFe2þ

of 0.82–0.88 for garnet cores, which decreases to0.71–0.77 towards the garnet rim, indicating progradegarnet growth (Fig. 4). Furthermore, both garnettypes are characterized by low grossular components,but nevertheless especially Grt2 shows a zone with aslight increase in XGrs near the rim that may beattributed to a decompressional garnet overgrowthduring increasing temperatures.The major element zoning patterns are compara-

ble to those of the trace elements; the latter wereanalysed to check for closed-system behaviour ofgarnet porphyroblasts during growth and exhuma-tion. In garnet, most LREEs are below the detec-tion limit, while heavy rare earth elements (HREEs)are relatively enriched. A common feature of allgarnet grains (Grt1 and Grt2) is a high degree ofHREE enrichment in the garnet core that decreasestowards the garnet rim, thereby excluding signifi-cant diffusion subsequent to garnet growth (Fig. 5).Given that both HREEs and MnO are prefer-entially fractionated in the porphyroblastic Grt1cores, the Grt2 grains, which started growing subse-quent to Grt1, have lower concentrations of theseelements.The foliation of the schist is mainly defined by talc,

the most abundant matrix mineral (~31 vol.%, meanXMg = 0.90), while quartz (~4 vol.%) and phengiticwhite mica (~3 vol.%, mean Si [per formula unit:p.f.u.] is 3.5) are less abundant. Matrix chloritoid(~16 vol.%, mean XFe2þ = 0.61) occurs as porphyro-blasts or forms aggregates; elongate minerals arecommonly oriented parallel to the foliation. Chlorite(~9 vol.%, mean XFe2þ = 0.25) occurs as a secondaryphase replacing garnet rims and chloritoid. Accessoryrutile (~1 vol.%) occurs as tiny round grains both inthe matrix and as inclusions in garnet, phengite andchloritoid.

Metasedimentary rocks: quartz-mica schists (10-02, 10-22)

Metasedimentary rocks of the central MakbalComplex are quartz-dominated and occur either as



relatively pure meta-quartzites or as quartz-micaschists. Two samples (10-02 & 10-22) chosen forthis study were collected at the central Makballocality. The metasedimentary rocks lack index min-erals such as (Fe,Mg)-carpholite, glaucophane, jade-ite or coesite that would have indicated (U)HPconditions. Both samples are quartz-dominated

(10-02: ~84.5 vol.%; 10-22: ~77 vol.%) with minoramounts of white mica (mainly phengite and veryrare paragonite; 10-02: ~9.2 vol.%; 10-22: ~7.8vol.%) that define a foliation. Phengite of sample10-02 has a mean Si [p.f.u.] of 3.51 and sample10-22 has a mean Si [p.f.u.] of 3.48. Sample 10-02contains in addition chlorite (~4.5 vol.%, mean

(a)

(c)

(b)

(d)

Fig. 2. (a) Field occurrence of a mafic Grt-Cld-Tlc schist. (b) Dextral shearing recorded by a garnet porphyroblast in a maficGrt-Cld-Tlc schist from the Makbal Complex. (c, d) Thin-section photographs of the Grt-Cld-Tlc schist investigated in this study(sample 10-16; crossed polars in (d)). Grt1 contains inclusions of chloritoid, talc, quartz, rutile and coesite, whereas Grt2 containsonly tiny rutile inclusions.



XFe2þ = 0.24) as matrix mineral or surroundingphengite and accessory apatite. Sample 10-22 addi-tionally contains porphyroblastic, strongly corroded(replaced by chlorite and/or paragonite) chloritoid(~3.5 vol.%, mean XFe2þ = 0.51) and secondarychlorite (~9.5 vol.%, mean XFe2þ = 0.22), whichoccurs in the matrix (mean XFe2þ = 0.22), replacingchloritoid (mean XFe2þ = 0.23) and surroundingphengite. In addition, besides accessory tourmaline(dravite–schorl), rutile, titanite and very rare zirconmay occur. The former presence of garnet insample 10-22 is indicated by a pseudomorphicintergrowth of chlorite, paragonite and quartz(Fig. 3d).

Mafic boudins

The petrography and petrology of the ‘eclogites s.l.’and the ‘glaucophanite’ (glaucophane-garnet-ompha-

cite-bearing rock) occurring as tectonic lenses in themetasedimentary host rocks are described in detail byMeyer et al. (2013) and thus only a short descriptionis given here.The main constituent of the ‘glaucophanite’ is

glaucophane (~42 vol.%) besides epidote-group min-erals (mainly clinozoisite, ~20 vol.%), garnet(~14 vol.%), quartz (~10 vol.%), sodic-calcic and rarecalcic amphibole (prograde relict as inclusion in gar-net, retrograde conversion rims around glaucophane,~8 vol.%). Minor and accessory minerals are phengit-ic white mica (~4 vol.%), omphacite (~1 vol.%),titanite (~1 vol.%), apatite and rutile (<1 vol.%). Themineral assemblages of the eclogites (s.l.) display anintense retrograde conversion under epidote-amphib-olite and/or greenschist facies conditions. The majorconstituents are amphibole (calcic and subordinatesodic-calcic; 26–48 vol.%), epidote-group minerals(mainly clinozoisite; 20–35 vol.%), garnet (5–19

(a) (b)

(c) (d)

Fig. 3. (a, b) Thin-section photographs (b: crossed polars) of the coesite inclusion in garnet of sample 10-16 (mafic Grt-Cld-Tlcschist). The coesite relict is surrounded by palisade quartz. (c) Raman spectrum of the coesite inclusion shown in (a) and (b).(d) Thin-section photograph of a poly-mineral aggregate (paragonite+chlorite+quartz+tourmaline) forming a garnet pseudomorph(sample 10-22, metasedimentary rock).



vol.%) and chlorite (0.2–20 vol.%). Minor constitu-ents are paragonite (2–7 vol.%), quartz (3–5 vol.%),rutile (2–5 vol.%) and titanite (0.3–3 vol.%) whileplagioclase, talc and apatite are accessories.

WHOLE ROCK GEOCHEMICAL DATA

UHP Grt-Cld-Tlc schist (sample 10-16)

The major and trace element concentrations for thissample are listed in Table 1. Remarkable are the highMgO and total Fe2O3 contents of 14 wt% (Mg-num-ber 44) and 20 wt%, respectively, at a low SiO2 of 43wt%. The REE pattern of the Grt-Cld-Tlc schist isshown in Fig. 6a and compared with other maficrocks (shaded region; data of other samples from thecentral Makbal Complex shown for comparison aregiven in Table S1) and another Grt-Cld-Tlc schist(sample 12-52, no coesite relicts, garnet compositionsimilar to that of Grt2 of sample 10-16) from thecentral Makbal Complex. The comparison of bothREE and trace element patterns shows a distinctivelydifferent pattern for sample 10-16: The LREE patternof sample 10-16 shows a negative slope with(La/Sm)N = 2.48, while the HREE elements arerelatively enriched compared with the MREE with

(Sm/Lu)N = 0.69 (Fig. 6a). The enrichment in HREEcorrelates with the high amount of garnet (~35vol.%), a major host of HREE (especially in Grt1,see Fig. 5) in the Grt-Cld-Tlc schist (sample 10-16).However, patterns of the high field strength elements(HFSE) and selected other elements are comparableto other samples. Some of the immobile HFSE (Th,Nb, Ta, Zr, Hf) have similar patterns for both Grt-Cld-Tlc schists (10-16 & 12-52), suggesting an identi-cal protolith for both samples (Fig. 6c). With theexception of sample 10-16, all other mafic rocks fromMakbal show a negative Nb-Ta anomaly in a primi-tive mantle (PRIMA)-normalized trace element dia-gram (Fig. 6c).

Metasedimentary rocks

The metasedimentary samples 10-02 and 10-22 havehigh SiO2 contents of ~92 and 87 wt%, respectively,corresponding to a high modal amount of quartz. Lowcontents of Al2O3 (5.5–7.3 wt%), MgO (1.25–2.67wt%) and total Fe2O3 (1.1–2.1 wt%) indicate a subor-dinate clay component. Furthermore, CaO contentsare 0.25 wt% for both samples, K2O concentrationsrange from 0.84 to 1.08 wt% and Na2O is below thedetection limit.

(a) (b)

Fig. 4. X-ray maps of Mn and Mg of thegarnet porphyroblast shown in Fig. 2d.Both Mn and Mg show an undisturbedprograde zoning pattern. The position ofthe cross-section is indicated on both maps.Spessartine and grossular are minorcomponents (right y-axis), but especially theXSps pattern shows a clear growth feature(‘bell-shaped’ pattern).



The chondrite-normalized REE patterns (Fig. 6b)of the samples are similar to that of the chemicallyevolved upper continental crust (UCC, Rudnick &Gao, 2003), albeit at different degrees of enrichmentdue to REE dilution by quartz. The PRIMA-normal-ized trace element patterns are similar for both sam-ples, even the fluid mobile large-ion lithophileelements (LILE) show uniform trends (Fig. S1).Major and trace element data of samples shown forcomparison (other metasedimentary rocks from theMakbal complex and passive margin sedimentaryrocks from McLennan et al., 1990) are summarizedin Table S1.

THERMOBAROMETRY AND THERMODYNAMICMODELLING: RECONSTRUCTION OF THEMETAMORPHIC EVOLUTION

Conventional thermobarometry

The application of conventional cation-exchangethermobarometry to the Grt-Cld-Tlc schist and

metasedimentary rocks is limited by the non-criticalmineral assemblages in these rock types. Thequartz-dominated metasedimentary rocks do notcontain assemblages suitable for this technique. TheGrt-Cld-Tlc schist contains both garnet andphengitic white mica, thus the garnet-phengite ther-mometer (Green & Hellman, 1982) was applied.Calculated temperatures range between 607 and635 °C for 2.7 GPa and 612 to 640 °C for2.8 GPa. The pressure values were chosen becausephengite is assumed to be a peak mineral and thepressure peak must have crossed the quartz–coesitetransition because relict coesite occurs as inclusionin garnet (Figs 2c,d & 3a–c) in accordance with thepressure estimate of Tagiri et al. (2010) of 2.8 GPa(at ~560 °C).

(a)

(b)

Fig. 5. Chondrite-normalized (Sun & McDonough, 1989) REEconcentrations in garnet porphyroblasts of the Grt-Cld-Tlcschist (10-16): (a) Higher concentration especially of HREE inGrt1-cores and the decrease in HREE from core to rim. (b)The same trend is obvious for Grt2-grains, even though HREEconcentrations in the garnet cores are lower.

Table 1. Whole rock major and trace element composition.Major elements are given in wt%, trace elements in ppm.

Sample 10-16 10-02 10-22

Rock type Grt-Cld-Tlc schist Metasedimentary rock Metasedimentary rock

Coordinates 42°46.2930 42°46.3620 42°46.2090

72°04.3410 72°04.2220 72°04.5080

SiO2 43.27 91.72 87.20

TiO2 0.86 0.09 0.14

Al2O3 17.18 5.50 7.28

Fe2O3 (total) 20.02 1.11 2.07

MnO 0.08 b.d.l. 0.01

MgO 14.25 1.25 2.67

CaO 0.53 0.25 0.25

Na2O b.d.l. b.d.l. b.d.l.

K2O 0.31 1.08 0.84

P2O5 b.d.l. 0.07 0.17

LOI 2.53 0.85 1.43

Total 99.03 101.92 102.06

Sc 49.2 3.58 3.59

Ti n.a. 832 1152

V 306 28.9 25.7

Cr 48.4 23.0 23.7

Mn 653 25.0 144

Co 34.8 4.71 6.51

Ni 84.3 13.5 13.4

Rb 4.43 21.9 14.9

Sr 3.12 7.20 7.88

Y 23.4 5.66 29.0

Zr 106 210 362

Nb 5.93 3.23 5.61

Cs n.a. 0.74 0.44

Ba 49.1 136 107

La 5.62 10.2 27.0

Ce 11.6 19.4 47.7

Pr 1.39 2.31 5.41

Nd 6.18 9.02 20.8

Sm 1.47 1.63 3.62

Eu 0.42 0.28 0.76

Gd 2.33 1.35 3.41

Tb 0.56 0.21 0.64

Dy 4.06 1.16 4.57

Ho 0.83 0.22 1.02

Er 2.53 0.60 3.05

Tm 0.35 0.10 0.46

Yb 2.42 0.71 2.97

Lu 0.35 0.12 0.45

Hf 2.72 5.37 9.20

Ta 0.49 0.28 0.54

Pb 2.68 2.72 2.79

Th 4.10 3.34 7.09

U 0.67 0.51 0.87

b.d.l., below detection limit; n.a., not analyzed.



THERMOCALC avPT

Metamorphic temperatures and pressures of the Grt-Cld-Tlc schist were determined using the programTHERMOCALC (version 3.3.6, using an updated version(2003) of the internally consistent Holland & Powell,1998 data set) in the average P–T mode (Powell &Holland, 1994) for the assemblage garnet+chlori-toid+phengite+talc+quartz with pure water in excess.Several calculations using coexisting mineral compo-sitions were conducted. Chloritoid is the onlyprimary inclusion in garnet rims, thus both chloritoidinclusions in garnet rims and matrix chloritoid wereused for avPT calculations, all other minerals arematrix minerals (Table S2). Irrespective of the tex-tural position of chloritoid, both equilibrium temper-

atures and pressures are relatively homogeneousranging between 531 � 14 to 541 � 11 °C and2.59 � 0.19 to 2.74 � 0.25 GPa (Table 2). Garnetgrains with a weak Ca increase near the rim indicateeven slightly higher P–T conditions, but equilibriumconditions are doubtful as indicated by high SigFitvalues.In general, prograde metamorphism may be recon-

structed by equilibrium phases included within garnetcores, but as mentioned above, the Grt-Cld-Tlc schistcontains predominantly chloritoid and rutile inclu-sions in garnet core and inner mantle sections; otherminerals like talc or chlorite are rare. Thus, the lownumber of minerals did not allow for a P–T estimateusing avPT.

Thermodynamic modelling

The metamorphic (P–T) evolution of both the Grt-Cld-Tlc schist (10-16) and the two metasedimentarysamples (10-02, 10-22) was modelled using the pro-gram PERPLEX (version 6.6.8, Connolly, 1990, 2005)and the internally consistent thermodynamic databaseof Holland & Powell (1998, with update of 2002).The accuracy and precision of any forward model-

ling are determined by uncertainties within the ther-modynamic database, the choice of activity modelsand considered thermodynamic components(elements either not measured or not considered inthe model system) and the degree of equilibriumattained between minerals. Despite these possibleuncertainties, pseudosection modelling is a powerfultool to reconstruct not only single points in P–Tspace (as for example with avPT or conventionalthermobarometry), but the P–T path of a sample bylocating stability fields of mineral assemblages andmineral reactions between stability fields. The inter-section of compositional isopleths of different miner-als and/or modal amounts of minerals withinstability fields allow for a more precise P–T recon-struction.The P–T pseudosections for the Grt-Cld-Tlc schist

were calculated in the MnCKFMASHT (MnO-CaO-K2O-FeO-MgO-Al2O3-SiO2-H2O-TiO2) model systemwith activity models for garnet (Holland & Powell,1998), chloritoid (White et al., 2000), white mica(Holland et al., 1998), talc (ideal) and chlorite (Hol-land et al., 1998). For the metasedimentary rocks,the CKFMASHT (CaO-K2O-FeO-MgO-Al2O3-SiO2-H2O-TiO2, sample 10-02) and MnCKFMASHT(sample 10-22) model systems were chosen with solu-tion models for phengite, chlorite, chloritoid and gar-net (Holland & Powell, 1998; Holland et al., 1998;White et al., 2000). Ferric iron was not consideredfor the two rock types, because the only major hostof Fe3+ in the Grt-Cld-Tlc schist is garnet, but allanalysed garnet has a very low Fe3+ content (Fe3+

recalculated using the method of Droop, 1987) andthe metasedimentary rocks do not contain minerals

(a)

(b)

(c)

Fig. 6. Chondrite-normalized whole rock REE and PRIMA-normalized (Sun & McDonough, 1989) trace elementconcentrations of (a, c) mafic rocks and (b) metasedimentaryrocks. Data for N-MORB, E-MORB and OIB are from Sun &McDonough (1989), UCC from Rudnick & Gao (2003). Insubfigure (a) and (c), besides the Grt-Cld-Tlc schist (10-16),other mafic rocks and one additional Grt-Cld-Tlc schist(12-52) from the Makbal Complex are shown for comparison(see Table S1), in (b) additional metasedimentary rocks(quartzites with different amounts of white mica and differentaccessory minerals) from the Makbal Complex are shownbesides data for metasedimentary rocks from McLennan et al.(1990) (see Table S1).



incorporating major amounts of Fe3+ (calculatedconcentrations of ferric iron in chloritoid and/orchlorite are very low, see Table 3). Given that bothrock types contain hydrous minerals like pheng-ite � chloritoid � talc that are believed to have beenstable during peak metamorphic conditions, andsulphides and graphite are absent, a fluid phase com-posed of pure H2O was considered to be in excess.

The appropriate effective bulk rock compositions(EBC) for the two metasedimentary samples werededuced from whole rock XRF results by correctingthe concentration of CaO for the amount containedin apatite, because P2O5 was not considered for mod-elling (see Table 4b). This simple approach is legiti-mate because both samples contain neither zonedporphyroblasts (i.e. no fractionation effects need tobe considered) nor heterogeneous domains on thinsection scale pointing at various equilibrium domainsrather than an overall equilibrium.

However, the occurrence of large, zoned garnetporphyroblasts in the Grt-Cld-Tlc schist (10-16)required the calculation of an EBC for this rock type.In this case, the calculation of an appropriate EBCfor different steps of garnet growth (and thus frac-tionation of elements in garnet, which are ‘removed’from the available bulk composition for subsequentmineral growth) is essential to obtain correct P–Tconditions of equilibration. Several methods havebeen proposed to calculate EBCs related to garnetgrowth (e.g. Spear et al., 1990; St€uwe, 1997; Marmoet al., 2002; Evans, 2004); in this study, the methoddescribed by Evans (2004) and Gaidies et al. (2006)was used (for more details on the application of thismethod, refer to Meyer et al., 2013). The initial EBCprior to garnet growth (i.e. M = 0.00, Table 4a) cal-culated from the modal amounts of minerals in thinsection (obtained from point-counting of several thinsections of the sample) and their average chemicalcomposition measured by EMPA (representativeanalyses are given in Table 3) is very similar to theresult from whole rock XRF analysis (Table 4a).

Mafic Grt-Cld-Tlc schist (10-16)

A pseudosection calculated for the initial EBC(M = 0.00) is shown in Fig. 7. Several minerals(e.g. lawsonite, zoisite, kyanite) occur in the modelledpseudosection, but not in thin section. This may beascribed either to the low modal amount (≤1 vol.%

in the relevant P–T regions) or to a retrograde con-version. The modal amount of lawsonite at the startof the P–T path shown in Fig. 7 is ~1 vol.% (Fig. 8),and lawsonite reacts out when the prograde P–T pathcrosses the lawsonite-out reaction at ~480 °C and2.55 GPa.Garnet is formed relatively late (garnet-in-reaction

is highlighted in Fig. 7), thus a prograde P–T pathdeduced from garnet composition and garnet inclu-sions cannot be used to reconstruct lower P–T condi-tions. Chloritoid inclusions in Grt1 are characterizedby a decrease of XFe2þ from 0.79 to 0.69 from core torim. Thus, both XFe2þ isopleths of chloritoid and raretalc inclusions in garnet with a maximum XMg of 0.92allow for a first approximation of P–T conditionsduring prograde garnet growth (see isopleths and darkblue shaded area in Fig. 7). During garnet outer-mantle to rim growth, the P–T path is located in thecoesite-stability field consistent with the observed coe-site-inclusion in one of the garnet porphyroblasts(Figs 2d & 3a–c). Other minerals occurring as inclu-sions in garnet do not show a systematic compositionalvariation during changing P–T conditions; therefore,they are not useful for a P–T path reconstruction.To assess a more detailed prograde to peak P–T

path, garnet isopleth thermobarometry was used. Thismethod requires the calculation of garnet isopleths fora changing EBC and is based on the assumption ofequilibrium between all phases (e.g. Spear & Selver-stone, 1983; Vance & Holland, 1993; Evans, 2004;Gaidies et al., 2006). Pseudosections were calculatedfor several steps of garnet growth (increasing valuesfor M, which stands for the amount of garnet alreadyformed) and corresponding successive fractionation ofelements from the ‘available’ bulk composition intogarnet (Table 4a). Intersections of garnet isopleths ofall analysed garnet grains define a P–T loop startingat ~480 °C at ~2.5 GPa reaching peak pressures of~2.85 GPa at 525 °C and peak temperatures of~580 °C at 2.4 GPa. This P–T path deduced from gar-net composition is in accordance with the compositionof matrix chloritoid with XFe2þ ranging between 0.56and 0.64 (light blue area in Fig. 7). Furthermore, thecomposition of zoned phengite (core: Si of 3.54[p.f.u.], rim: Si of 3.52 [p.f.u.]) fits well that of themodelled phengite after passing peak-pressure condi-tions, and is in accordance with the measured XMg of0.90 in talc (isopleths of XFe2þ in chloritoid, XMg intalc and Si [p.f.u.] in phengite are shown in Fig. 7).

Table 2. Summary of avPT results (sample 10-16), N gives the number of independent reactions.

Sample Equilibrium assemblage T (°C) P (GPa) DT (°C) DP (GPa) SigFit N

10-16-14 (Grt1+Cld incl) Grt + Cld + Ph + Tlc + Qz + H2O 532 2.68 16 0.29 1.54 5

10-16-14 (Grt1+matrixCld) Grt + Cld + Ph + Tlc + Qz + H2O 531 2.74 14 0.25 1.32 5







Table

3.Representativemineralchem

icaldata

oftheGrt-C

ld-Tlc

schist(10-16)andmetasedim

entary

rocks(10-02,10-22)from

thecentralMakbalComplex.

Sample

10-16

10-02

10-22

Rock

type

Grt-C

ld-Tlc

schist

Qz-micaschist

Qz-micaschist

Mineral

Grt

Grt

Grt

Grt

Grt

Grt

Grt

Tlc

Ph

Cld

Cld

Chl

Ph

Pg

Chl

Ph

Pg

Cld

Chl

Tur

Position

Grt-1-c

Grt1-m

1Grt1-m

2Grt1-r

Grt2-c

Grt2-m

Grt2-r

Grt1-i

SiO

236.35

36.10

36.49

37.72

37.53

37.76

38.15

61.75

55.13

24.65

24.99

28.42

55.84

50.41

28.61

55.72

49.18

25.77

28.53

36.92

TiO

20.09

0.05

0.03

0.00

0.00

0.06

0.00

0.00

0.23

0.00

0.00

0.03

0.20

0.04

0.08

0.22

0.00

0.00

0.00

0.74

Al 2O

321.52

21.76

22.24

22.38

21.41

21.67

22.21

0.17

25.84

40.79

42.48

20.80

25.01

39.32

20.82

25.32

38.99

42.17

21.90

32.83

Cr 2O

30.01

0.00

0.00

0.05

0.02

0.01

0.02

0.01

0.05

0.00

0.04

0.03

0.05

0.00

0.00

0.00

0.01

0.07

0.01

0.00

Fe 2O

3n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

FeO

(total)

37.18

37.59

34.46

33.60

35.94

35.17

32.97

5.41

1.89

23.07

17.72

14.79

0.99

0.15

12.36

1.17

0.27

15.12

12.42

3.58

MnO

1.15

0.67

0.18

0.14

0.19

0.14

0.06

0.00

0.00

0.06

0.01

0.01

0.00

0.00

0.04

0.00

0.00

0.22

0.06

0.03

MgO

3.14

3.59

5.75

7.07

4.60

5.38

7.23

27.94

4.19

3.78

7.10

23.96

4.91

0.14

25.03

4.75

0.13

8.62

24.40

9.09

CaO

0.98

0.77

0.75

0.35

0.63

0.27

0.18

0.04

0.03

0.02

0.00

0.06

0.03

0.23

0.07

0.04

0.25

0.02

0.01

1.67

Na2O

0.14

0.03

0.04

0.10

0.06

0.07

0.08

0.00

0.29

0.00

0.00

0.02

0.27

7.17

0.01

0.29

6.46

0.02

0.03

1.50

K2O

0.01

0.00

0.01

0.00

0.00

0.01

0.00

0.00

10.09

0.00

0.00

0.00

10.38

0.82

0.02

10.08

0.78

0.01

0.00

0.01

Total

100.56

100.56

99.96

101.41

100.38

100.55

100.90

95.33

97.75

92.36

92.34

88.11

97.68

98.28

87.04

97.59

96.07

92.02

87.36

86.36

Si

2.92

2.89

2.89

2.92

2.99

2.99

2.96

7.99

3.55

2.02

1.99

5.60

3.59

3.11

5.63

3.58

3.10

2.04

5.58

6.40

Ti

0.01

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.01

0.00

0.00

0.00

0.01

0.00

0.01

0.01

0.00

0.00

0.00

0.10

Al

2.04

2.05

2.07

2.04

2.01

2.02

2.03

0.03

1.96

3.94

3.99

4.84

1.89

2.86

4.84

1.92

2.89

3.93

5.07

6.71

Cr

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.01

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

Fe3

+0.11

0.16

0.15

0.11

0.02

0.00

0.04

0.00

n.d.

0.00

0.00

0.04

n.d.

n.d.

0.08

n.d.

n.d.

0.07

0.13

0.00

Fe2

+2.39

2.35

2.13

2.07

2.37

2.32

2.11

0.59

0.10

1.58

1.18

2.39

0.05

0.01

1.96

0.06

0.01

0.93

1.90

0.52

Mn

0.08

0.05

0.01

0.01

0.01

0.01

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.01

0.00

0.00

0.01

0.01

0.00

Mg

0.38

0.43

0.68

0.82

0.55

0.63

0.84

5.39

0.40

0.46

0.84

7.04

0.47

0.01

7.35

0.45

0.01

1.01

7.12

2.35

Ca

0.08

0.07

0.06

0.03

0.05

0.02

0.01

0.01

0.00

0.00

0.00

0.01

0.00

0.01

0.01

0.00

0.02

0.00

0.00

0.31

Na

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

0.00

0.04

0.00

0.00

0.01

0.03

0.86

0.01

0.04

0.79

0.00

0.02

0.50

Kn.d.

n.d.

n.d.

n.d.

n.d.

n.d.

n.d.

0.00

0.83

0.00

0.00

0.00

0.85

0.06

0.01

0.83

0.06

0.00

0.00

0.00

Sum

cation

88

88

88

814

78

820

77

36

77

836

17

O12

12

12

12

12

12

12

22

11

12

12

28

11

11

28

11

11

12

28

31

XFe

0.86

0.85

0.76

0.72

0.81

0.79

0.72

0.10

0.77

0.58

0.25

0.21

0.21

XAlm

0.82

0.81

0.74

0.71

0.79

0.78

0.71

XGrs

0.03

0.02

0.02

0.01

0.02

0.01

0.00

XPrp

0.13

0.15

0.23

0.28

0.18

0.21

0.28

XSps

0.03

0.02

0.00

0.00

0.00

0.00

0.00

c,core;i,inclusion;m,mantle;

n.d.,notdetermined;r,rim.



The retrograde evolution of this sample is difficultto constrain, because the only mineral grown duringretrogression is chlorite replacing both garnet andchloritoid. Even though XFe2þ of chlorite varies atchanging P–T conditions, no further reconstruction

of the retrograde P–T path is possible, becausechlorite is modelled to grow rapidly after crossing thechlorite-in-reaction.The pseudosection of the Grt-Cld-Tlc schist is also

contoured for the modal amount of garnet, indicating

Table 4. Summary of chemical composition used as input for PerpleX. (a) Major element composition of sample 10-16: XRF andEBC compositions are given. Note the similar composition of EBC II (WR-EBC minus garnet = matrix composition) and the EBCfor M = 0.35 (garnet rim growth) calculated for successive incorporation of elements in garnet and removal from the ‘available’bulk composition. (b) Major element composition of the two metasedimentary samples (10-02 and 10-22) used for modelling.

SiO2 TiO2 Al2O3 FeO (total) MnO MgO CaO Na2O K2O Total

(a) Grt-Cld-Tlc schist (sample 10-16)

XRF analysis [wt%] normalized to 100% anhydrous

XRF (WR) 45.79 0.91 18.18 19.06 0.08 15.08 0.56 b.d.l. 0.33 100

Effective bulk composition calculated from mineral modes and microprobe analyses [wt%]

EBC (WR) 46.53 1.40 17.65 19.36 0.12 14.37 0.22 0.04 0.32 100

EBC II (WR without garnet = matrix) 51.95 2.18 15.45 9.99 0.01 19.87 0.02 0.03 0.50 100

Effective bulk composition for modelling (MnCKFMASHT) [wt%] – modification of EBC (WR) according to successive garnet growth (M = 0.00–0.35)M = 0.00 46.55 1.40 17.66 19.37 0.12 14.37 0.22 n.c. 0.32 100

0.00 < M < 0.35: calculated in several steps

M = 0.35 51.62 2.16 15.56 10.62 0.00 19.52 0.04 n.c. 0.49 100

(b) Metasedimentary rocks

XRF analysis [wt%] normalized to 100% anhydrous, corrected for apatite-free composition

10-02 90.99 0.09 5.46 0.99 b.d.l. 1.24 0.16 b.d.l. 1.07 100

10-22 86.98 0.14 7.26 1.86 0.01 2.66 0.25 b.d.l. 0.84 100

b.d.l., below detection limit; n.c., not considered.

Fig. 7. P–T pseudosection calculated forthe Grt-Cld-Tlc schist (10-16) prior togarnet growth (i.e. M = 0.00) in the modelsystem MnCKFMASHT. Simplifiedreactions are indicated by thicker lines.Besides the modal amount of garnet,isopleths for Si in phengite, XMg in talc andXFe2þ in chloritoid are shown. The blueshaded area covering the range of XFe2þbetween 0.79 and 0.69 corresponds tochloritoid inclusions in garnet and the lightblue area (XFe2þ between 0.64 and 0.56)represents matrix chloritoid compositionsThe P–T path is deduced from garnetisopleth thermobarometry and thecomposition of garnet inclusions andmatrix minerals. Solid solution series ofminerals are indicated by capital letters.



a prograde rapid garnet increase close to the chlorite-out reaction under HP conditions (Fig. 7). Thisobservation is in accordance with rapid garnetgrowth during subduction-related H2O release asobserved by Baxter & Caddick (2013). To better con-strain the changes in mineral modes and water con-tent, their modal amounts (mineral modes in vol.%)and weight amount (XH2O in wt%), respectively, weremodelled along the prograde P–T path shown inFig. 7. Modelled mineral modes (Fig. 8) correlatewell with those observed in thin section with theexception of chlorite, which is not stable at the endof the P–T loop, and talc, the modal amount ofwhich is overestimated. In addition, modelling pre-dicts that the rock releases a total of 5.75 wt% H2Oalong the P–T loop during garnet growth. Thisamount lies within the dehydration capacity of well-hydrated MORB, representative of oceanic crust(3.3–5.9 wt%) and well above that of pelite (1.8–3.1wt%, Baxter & Caddick, 2013). Two major dehydra-tion events correlate with major garnet production(Fig. 8) and confirm the correlation between garnetgrowth and progressive dehydration of hydrous min-erals in subduction zones (e.g. Dragovic et al., 2012;Baxter & Caddick, 2013).

Metasedimentary rocks (10-02, 10-22)

The reconstruction of the P–T path or peak meta-morphic P–T conditions of the metasedimentary

rocks is not as precise as for the Grt-Cld-Tlc schistbecause zoned minerals recording part of the P–Tevolution (like garnet) are absent. Thus, maximumpressures were estimated from the modelled and mea-sured Si content in white mica for both samples andfor sample 10-22; the composition of chloritoid wasadditionally applied to estimate P–T conditions.A P–T pseudosection calculated for the metasedimen-

tary sample 10-02 (Fig. 9a) suggests that this samplemight have experienced HP conditions. As mentionedabove, critical minerals that would prove HP or evenUHP conditions (e.g. coesite) are absent. But the veryhigh Si content of elongate (prismatic) phengite (maxSi of 3.60 [p.f.u.]) suggests at least HP conditions dur-ing peak metamorphism (Si isopleths of phengite inFig. 9a). The retrograde evolution is only representedby chlorite with XFe2þ ranging between 0.21 and 0.26corresponding to lower pressure conditions (XFe2þ

isopleths of chlorite in Fig. 9a). Pseudosection model-ling suggests the existence of up to 5 vol.% chloritoidover a wide P–T range, even though it is not observedin sample 10-02. However, the amount of chloritoiddecreases towards lower P–T conditions; therefore, itis likely that it was replaced by chlorite during retro-gression. This is in accordance with the petrographicalobservations. Additional minerals such as lawsonite,magnesio-carpholite, talc and kyanite are also sup-posed to be stable in the modelled P–T range.However, all of them are predicted to be present asminor (<5 vol.%) or accessory (<1 vol.%) phases.Results of pseudosection modelling of sample 10-

22 are similar to those of sample 10-02. Even thoughthis sample contains partially retrogressed chloritoidminerals, precise peak P–T conditions are impossibleto retrieve. As described for sample 10-02 above, thissample also exhibits very high Si contents in lepidob-lastic phengite minerals (max Si of 3.59 [p.f.u.])pointing to HP or even UHP conditions (Fig. 9b).Another peak metamorphic mineral (implying HPconditions) is chloritoid with a maximum XFe2þ of0.54. Furthermore, the occurrence of a poly-mineral(paragonite+chlorite+quartz+tourmaline) garnet pseu-domorph (Fig. 3d) implies that during the progrademetamorphic evolution, the garnet-in-reaction musthave been crossed. The post-peak metamorphic evo-lution is recorded in phengite with slightly lower Sicontents (Si <3.5 [p.f.u.]) that is intimately intergrownwith chlorite (mean XFe2þ = 0.24) replacing chloritoidand thus pointing to a retrograde growth of thesesubidioblastic minerals. Rare paragonite probablyformed under lower pressure conditions with a Si of3.10 [p.f.u.].

Mafic rocks

In this study, no pseudosection modelling and/or gar-net isopleth thermobarometry was conducted for themafic boudins, as it had been done already in a previ-ous study (Meyer et al., 2013).

Fig. 8. Changes in modal proportions of minerals (or solutionphases) and the amount of water. Modelling was conductedalong the P–T path shown in Fig. 7 (circles in Fig. 7correspond to grey shaded lines in this figure). Majordehydration events along the P–T loop correlate with garnetgrowth. Solid solution series of minerals are indicated bycapital letters.



DISCUSSION

Sm-Nd dating of the Grt-Cld-Tlc schist (sample 10-16)

The Sm-Nd isotope data of the whole rock, a garnetseparate and matrix material (whole rock materialfrom which garnet was separated) are listed inTable 5 and plotted in an isochron diagram inFig. 10. The low Sm/Nd ratio of the whole rockpowder indicates an LREE-enriched composition

with an unusually low eNd value of �11.3 at480 Ma. If this sample has been derived by mantlemelting and ocean basin development in the north-ern Tianshan at c. 510–530 Ma (Kr€oner et al.,2012), it would indicate a mantle source with aneNd of ~�11, which can be precluded as a meaning-ful value for the upper mantle (an unrealistically oldage of 2.5 Ga would give the mafic sample adepleted mantle-like value of +1). Melting of anenriched Archean lithospheric mantle source alsoappears as unlikely interpretation because the base-ment was formed during the Neoproterozoic, andthereafter, and evidence for Archean basement withthick mantle roots is missing. More plausible is theassumption that the Sm-Nd system in the samplewas overprinted during rock alteration and meta-morphism, implying that other LREE and similarlybehaving incompatible elements were also modified.The unusually low eNd value at 480 Ma indicatessample overprinting by very non-radiogenic Nd pos-sibly leached during plate subduction from sedimentwith ancient crustal residence times. The good agree-ment between the isotope dilution and ICP-MS dataprecludes an analytical artefact. In the isochrondiagram (Fig. 10), the whole rock and garnet data

(a) (b)

Fig. 9. P–T pseudosections for the metasedimentary rocks, simplified reactions are indicated by thicker lines: (a) Sample 10-02modelled in the CKFMASHT model system. The blue area indicates the range of Si in phengite, the darker blue area correspondsto the XFe2þ of retrograde chlorite probably replacing chloritoid. Chloritoid is not observable in thin section, which is inaccordance with the low modal amount up to 5 vol.% (isopleths are indicated by dashed lines). (b) Sample 10-22 modelled in theMnCKFMASHT model system. The garnet-in reaction is shown because of the preservation of garnet pseudomorphs. The bluearea indicates the range of Si in phengite; the area with lower values (Si of 3.38–3.31 [p.f.u.]) shows the compositional variation ofphengite replacing chloritoid. Furthermore, the purple area indicates the observed composition of chloritoid and the dark blue areacorresponds to the XFe2þ of retrograde chlorite, often intergrown with low-Si phengite.

Table 5. Sm-Nd isotopic data for Grt-Cld-Tlc schist (sample10-16) from the central Makbal Complex, northern Tianshan,Kazakhstan/Kyrgyzstan.

Sample Sm

(lg g�1)

Nd

(lg g�1)

147Sm/144Nd 143Nd/144Nd (m.) eNd

(480 Ma)

Whole rock 1.517 6.149 0.1491 0.511903 � 5 �11.3

Whole rock

matrixa1.52 6.362 0.1444 0.511819 � 5

Garnet 0.147 0.0717 1.2405 0.515302 � 26

aRock powder prepared from garnet-mantling material; 143Nd/144Nd normalized to146Nd/144Nd = 0.7219. External precision for 143Nd/144Nd is ~5 9 10�6 (2SD). Error for147Sm/144Nd 0.15% (2SD). The JNdi-1 reference material yielded during the course of

this study 143Nd/144Nd = 0.512103 � 3 (2SD, n = 8). m. = measured ratio, error is 2rmean. eNd calculated with parameter of Bouvier et al. (2008). Input errors (2SD):147Sm/144Nd = 0.15%; 143Nd/144Nd = 0.001% for whole rock and 0.005% for garnet

analysis.



suggest an average age for garnet growth of475 � 4 Ma, which is indistinguishable within errorsfrom the 470 � 2.5 Ma Lu-Hf garnet age foranother mafic rock sample from Makbal (Rojas-Agramonte et al., 2013). The data point of the sam-ple matrix plots a little below the garnet–whole rockisochron, indicating that the entire matrix had notachieved isotopic equilibrium with the garnet popu-lation. Age regression analysis of matrix, whole rockand garnet data give an errochron age of c. 480 Ma –not much different from the whole rock–garnet age.The average garnet growth age of 475 Ma inferredfrom the Sm-Nd system agrees with other publishedchronological data for the Makbal region thatbracket metamorphism from c. 500 to 470 Ma. Wenote that these data were obtained by different iso-tope systems and minerals with different closuretemperatures from that of the Sm-Nd system ingarnet. In addition, these samples all had distinctiveP–T–t evolutions interpreted to have developed in adynamic subduction–exhumation process. For themillimetre-sized garnet grains processed in this study,we suggest that the Sm-Nd age represents the peaktemperature range of <600 to 650 °C (Rojas-Agra-monte et al., 2013; this study). Due to a smallerionic radius, Mn2+ is supposed to diffuse faster thanSm and Nd, but given that its prograde growth zon-ing is well preserved (Fig. 4), it is to be expectedthat the closure temperature of the Sm-Nd systemwas not exceeded. Thus, it is believed that agesobtained from the Sm-Nd geochronology representprograde garnet growth. Furthermore, the overallagreement of the Sm-Nd garnet age with the pub-lished ages for HP metamorphism indicates as wellthat the Sm-Nd isotope system represents peakmetamorphic conditions; however, the high-precisionisotope measurements reveal some degree of isotopicdisequilibrium between sample matrix and garnetduring garnet growth that maybe due to the strong

retrogression during exhumation of the sample. Iso-topic disequilibrium between matrix minerals andgarnet has been reported for example by Th€oni(2002). He considered this phenomenon to be impor-tant in mafic eclogites where the Sm-Nd system inmatrix minerals may close late after peak-pressureconditions and after closed system isotope evolutionhad been established in garnet. In such cases,matrix–garnet Sm-Nd ages would be unreliable. Weinterpret the matrix–garnet age of this study simi-larly and accept the whole rock–garnet age as geo-logically meaningful under the condition thatisotopic equilibrium had been established on thescale of whole rock and garnet during peak meta-morphic conditions.

Protoliths of (U)HP rocks

Constraining the protolith of oceanic mafic (U)HPmetamorphic rocks may be difficult due to significantmodifications of the major and trace elements byhigh- and/or low-T interaction with seawater and thesubsequent metamorphism. Particularly, the highlyfluid mobile LILE are likely to be mobilized duringsubduction (e.g. John et al., 2003). Therefore, themost reliable trace elements for identifying precursorrocks of high-grade metamorphic units are relativelyfluid-immobile trace elements like the REE and theHFSE (e.g. M€oller et al., 1995; John et al., 2004;Miller et al., 2007).

Grt-Cld-Tlc schist (10-16)

Previous studies dealing with Grt-Cld-Tlc schists fromMakbal suggested a pelite (Tagiri et al., 2010) or a lat-eritic clay (Bakirov et al., 2008) as potential protoliths,but did not provide any geochemical evidence for thisassumption. In our previous study (Meyer et al.,2013), the Grt-Cld-Tlc schist still was referred to as ametasedimentary rock. This classification, however,may need to be revised using the new geochemicaldata of this study: The high MgO, Fe2O3 and lowCaO and alkali element concentrations can best beexplained by a pervasively altered mafic igneous rockrather than a sedimentary rock. Especially, the lowconcentrations of the highly fluid-mobile elementssuch as K2O (0.31 wt%), Na2O (below detection limit)and CaO (0.53 wt%) and trace elements like Rb andSr indicate significant removal of these elements. Ithas been shown that high-T seafloor alteration maylead to MgO gain, and loss of CaO and alkali elementsas well as a minor SiO2 loss (e.g. Humphris & Thomp-son, 1978; Mottl & Holland, 1978; Seyfried & Mottl,1982). Such hydrothermal alteration is typically asso-ciated with oceanic basins rather than mafic dykeinjection into continental crust. The latter scenario hasbeen suggested by Kr€oner et al. (2012) and Rojas-Agramonte et al. (2013) as the protolith for maficrocks in the Aktyuz and marginal Makbal meta-

Fig. 10. Sm-Nd isochron diagram for whole rock (Wr), matrixand garnet (Grt) data of Grt-Cld-Tlc schist (sample 10-16).



morphic complexes. However, for Aktyuz, Kr€oneret al. (2012) were able to provide field evidence to sup-port their interpretation. Nevertheless, at Makbal, theprotolith of the Grt-Cld-Tlc schist was presumably anintensely metasomatized magmatic rock, eitherstrongly altered oceanic crust, which is in line with theinterpretation that other mafic boudins of the MakbalComplex (eclogites s.l. and ‘glaucophanite’, Meyeret al., 2013) represent oceanic crust less affected byhydrothermal alteration, or volcaniclastic materialfrom a magmatic arc that built on the oceanic crustprior to subduction. Besides the unusually low eNdvalue of �11.3, the latter interpretation is supportedby the strongly variable SiO2 contents (43–59.6 wt%)of the Grt-Cld-Tlc schist investigated in this study andanother Grt-Cld-Tlc schist from the Makbal Complex(sample 12-52). A comparison of these two Grt-Cld-Tlc schists shows that especially the LREEs andMREEs of sample 10-16 clearly deviate not only fromsample 12-52 but also from the magmatic patterns ofthe other mafic rocks from Makbal (shaded region)and no longer represent pristine igneous concentra-tions (Fig. 6a). This indicates a mobilization of theseelements that may have occurred either during hydro-thermal alteration of the ocean floor or during subduc-tion.

The lack of a Nb-Ta anomaly in sample 10-16 ismost likely caused by LREE mobilization (especiallyLa and Ce) and is not a primary feature (Fig. 6c).Even though the Nb-Ta anomalies of the othersamples shown for comparison in Fig. 6c are not

very pronounced, they suggest protoliths with a sig-nificant contribution from a subduction zone compo-nent as observed in arc basalts or incipient back-arcbasins (e.g. Gill, 1981; Pearce & Stern, 2006), giventhat in modern back-arc basins like the East ScotiaRidge (South Atlantic Ocean), Nb-Ta anomalies maynot be very pronounced (e.g. Fretzdorff et al., 2002;see Fig. S2). In addition, the Th/Yb v. Nb/Yb rela-tionship given on Fig. 11a shows that all mafic rocksfrom the Makbal Complex plot either in, or above,the oceanic basalt [mid ocean ridge basalts (MORB)– ocean island basalts (OIB)] array. Variations in theTh–Nb relationship are attributed to crustal involve-ment resulting either from crustal contamination orsubduction-related crustal recycling (Pearce, 2008).The variability in Th, which is mobile in a subduc-tion zone setting (Pearce, 2008) and is enriched in arcand intra-arc rift basalts, is responsible for the analy-ses plotting above the MORB-OIB array (Pearceet al., 2005). Basalts from back-arc basins maybehave transitionally, plotting in or above theMORB-OIB array depending on their arc proximity(Pearce, 2008). The mafic rocks from Makbalare characterized by Nb/Yb ratios of 1.76–4.24(E-MORB: 3.50), but compared with E-MORB(Th/Yb = 0.25) they all have higher Th/Yb ratios(0.33–2.48) (Fig. 11a). Therefore, we suggest thatthey were all derived from the same enriched mantlesource (similar melting depth, cf. Fig. 11b), but withdifferent (but still minor) proportions of subductioninput (Fig. 11a).

(a) (b)

Fig. 11. (a) Th/Yb v. Nb/Yb diagram of Pearce (2008) utilizing the Th-Nb ratio as a proxy for crustal input. Data of mafic rocksfrom the central Makbal Complex (black diamonds, sample 10-16: blue circle) plot in most cases above the oceanic MORB–OIBarray, suggesting different degrees of Th enrichment that is derived from subduction-related crustal recycling. Note that input of≤5% of a crustal component (modelling adapted from Pearce, 2008; Fig. 5d) results in a significant shift of basaltic compositionsabove the MORB-OIB array. (b) TiO2/Yb v. Nb/Yb diagram of Pearce (2008) utilizing the TiO2/Yb ratio as a proxy for meltingdepth. All mafic rocks plot in, or close above the MORB-array, indicating a similar, shallow melting depth.



Metasedimentary rocks (10-02, 10-22)

The high concentration of SiO2 points to a clasticsedimentary protolith that was presumably depositedon a continental shelf. In addition, detrital zircon inmetasedimentary rocks of the Makbal Complex withages ranging between 3300 and 1900 Ma (U-PbSHRIMP, Degtyarev et al., 2013) indicates a clasticsource of the metasedimentary rocks. A passive mar-gin setting is supported by several observations: themetasedimentary rocks are quartz-dominated withevolved major element compositions (high SiO2/Al2O3 ratios between 17 and 12). Furthermore, allsamples are characterized by negative Eu-anomalies(Eu/Eu* = EuN/(SmN*GdN); 0.57–0.65, Fig. 6b) andare enriched in incompatible elements, which allpoint to a mature continental source, possibly a pas-sive margin (McLennan et al., 1990). The REE-pat-terns of both samples show that especially the LREEare more enriched than the HREE (10-02:(La/Sm)N = 4.02, (La/Yb)N = 10.32; 10-22: (La/Sm)N = 4.81, (La/Yb)N = 6.51).

General subduction and exhumation mechanisms ofoceanic and continental crust

The occurrence of both HP and UHP eclogite faciesoceanic mafic rocks at the central Makbal Complexhosted by (U)HP metasedimentary rocks (e.g. Meyeret al., 2013; present study) and continental maficrocks at the marginal (see Fig. 1b) Makbal Complex(e.g. Rojas-Agramonte et al., 2013) requires a sub-duction and exhumation mechanism incorporatingboth oceanic and continental crust.

Reviewing more than 60 Phanerozoic HP-UHPterranes, Guillot et al. (2009) suggested three maintypes of subduction zones and associated exhumationmodels:

(1) Accretionary wedge (or prism) type subductionzones develop in front of intra-oceanic arcs or conti-nental arcs and are characterized by the stackingmainly of sedimentary rocks due to offscraping ofsediments either derived from the lower plate or byarc erosion (e.g. Cloos, 1982; Platt, 1986; Moore &Silver, 1987; Cloos & Shreve, 1988a). This modelincludes a syn-convergent tectonic exhumation pro-cess that suggests corner flow circulation in relativelysmall accretionary wedges and is thus generally validfor exhumation of shallowly to intermediately sub-ducted oceanic rocks (Cloos, 1982). Even though thismodel was later generalized to larger wedge zones toexplain the exhumation of HP rocks (e.g. Platt,1993), it does not sufficiently explain the exhumationof deeply subducted (up to UHP conditions) rocks,because the maximum depth (mainly controlled bythe geometry of the prism, see Ernst, 2005; Guillotet al., 2009) of an accretionary wedge in the present-day subduction settings is ~20 km (~0.6 GPa) and

may reach depths of 40–60 km in exceptional cases(Guillot et al., 2009).Another mechanism to subduct and exhume oce-

anic metabasites and metasedimentary rocks can bededuced from the common association of HP–UHProcks and serpentinite in the so-called (2) serpenti-nite-type subduction zones (serpentinite-type subduc-tion channels). Commonly, highly shearedserpentinites (mostly originated from subducted abys-sal peridotites and/or the hydrated mantle wedge)constitute the matrix and contain weakly deformedblocks (cm- to km-scale) of metabasites and subordi-nate (<10 vol.%) highly deformed metasedimentaryrocks (e.g. Agard et al., 2009; Guillot et al., 2009).The higher buoyancy of serpentinite is regarded asfacilitating the exhumation of the more dense eclog-ites in the subduction channel (e.g. Cloos, 1982).Even though this model is also applicable for deeplysubducted oceanic crust (down to the stability limitof serpentinite), it cannot explain occurrences of(U)HP oceanic eclogites without serpentinite as thesedimentary m�elanges of the Tianshan at Makbal orthe southwest Tianshan of China.(3) The continental-type subduction zone (alpine-

type subduction) involves the subduction and con-sumption of oceanic crust that is followed by thesubduction of a thinned continental margin that canbe subducted to HP–UHP depths of at least 100 km.Exhumation from such depths is attributed to thepositive buoyancy of surrounding quartzofeldspathicrocks. Several geodynamic processes like eduction,microplate rotation, crustal stacking, slab rollback-trench retreat, channel flow and diapirism are envi-sioned to have been or can be responsible for theexhumation of HP–UHP quartzofeldspathic terraneswith subordinate mafic and ultramafic rock (fordetails see Hacker & Gerya, 2013; Warren, 2013). Infact, most models dealing with buoyancy consider theoceanic crust as denser than the surrounding mantle,which requires additional, less dense material as atransport agent for exhumation. However, Chenet al. (2013) suggested that the high amount of low-density HP–UHP hydrous minerals (e.g. glauco-phane, lawsonite, paragonite, phengite, talc, chlori-toid) in low-T eclogites (which form the majority ofall exhumed eclogites) may result in a lower densityof these rocks (former upper part of the subductedoceanic crust) than the mantle up to depths of� 110 km along a cold subduction path with a geo-thermal gradient of 6 °C km�1.

Formation of the Makbal tectonic m�elange

The modelling of P–T paths in this study and ear-lier work (Meyer et al., 2013) shows that fragmentsof former oceanic crust found at Makbal were sub-ducted to different maximum depths (HP v. UHPconditions) following different P–T trajectories (the



term oceanic crust used in the following discussionrefers to both proposed protoliths for the Grt-Cld-Tlc schists irrespective of whether it was a volcanicflow or volcaniclastic material because they are bothpart of the subducted oceanic slab). The peak meta-morphic conditions of the metasedimentary hostrocks are less well constrained (see above), but it islikely that they reached at least HP, or even UHPconditions (e.g. Togonbaeva et al., 2009; Tagiriet al., 2010). Besides mafic blocks and boudins, Tag-iri et al. (2010) reported the occurrence of ‘exoticlenses’ (calcsilicate schist, biotite-phengite-garnetschist and winchite schist), which had experiencedmaximum pressures of ~0.8 GPa (Tagiri et al.,2010), significantly lower than those of the maficblocks, and also lower than those of the surround-ing metasedimentary rocks. Such an association oflithologies derived from different depths of the sub-duction zone is thought to be typical for tectonicm�elanges formed by tectonic juxtaposition duringsubduction and exhumation (e.g. Cloos, 1982;Shreve & Cloos, 1986; Cloos & Shreve, 1988b;Federico et al., 2007; Agard et al., 2009).

At Makbal, metabasic rocks (Meyer et al., 2013;sample 10-16, this study) and sedimentary rocks(samples 10-02, 10-22, this study) were together sub-ducted and experienced HP or UHP metamorphism.The exhumation of disrupted fragments of subductedoceanic crust in a hydrated most likely metasedimen-tary-dominated subduction channel can be facilitatedby both the large amount of light hydrous mineralsin the mafic rocks (e.g. Chen et al., 2013) and aforced channel flow (e.g. Gerya et al., 2002). Thisexhumation mechanism is similar to the serpentinite-type subduction channels of Guillot et al. (2009), butas serpentinites are missing at the Mabal m�elange,they were either not exhumed to shallow crustaldepths or the required buoyancy in the subductionchannel was entirely supplied by metasedimentaryrocks. The observed differences in the modelled P–Tpaths (HP v. UHP) may be attributed to exhumationfrom different depths in the subduction channel orslab, and/or differences in the subduction setting(change in slab angle, cf. Klemd et al., 2011).

Geodynamic formation of the Northern Tianshan andexhumation mechanism at Makbal

The current tectonic model for the Anrakhai complexand the Aktyuz metamorphic complex favours thatthe North Tianshan formed as a microcontinent sur-rounded by the Djalair (or Dzhalair)-Naiman basinto the north and the Kyrgyz-Terskey basin to thesouth during the Cambrian and Early Ordovician(e.g. Avdeev & Kovalev, 1989; Bakirov & Maksum-ova, 2001; Ghes’, 2008; Alexeiev et al., 2011; Kr€oneret al., 2012; Degtyarev et al., 2013; Rojas-Agramonteet al., 2013). Convergence and subduction of theDjalair-Naiman basin (a back-arc basin formed

through southwest-directed subduction of the Paleo-Asian Ocean under the northeastern margin of theNorth Tianshan terrane during the Late Neo-Proterozoic and Early Cambrian; Alexeiev et al.,2011; Kr€oner et al., 2013) towards the northeastunder the southwestern margin of the Anrakhai mi-crocontinent are suggested to have started in the LateCambrian (e.g. Alexeiev et al., 2011; Kr€oner et al.,2013). In the Early Ordovician (c. 475 Ma), the Djal-air-Naiman basin was closed and presumably due toslab pull, the previous passive margin of the NorthTianshan microcontinent was involved in northeast-directed subduction underneath the Anrakhai micro-continent (Kr€oner et al., 2012; Rojas-Agramonteet al., 2013). The depth of continental subduction(initiated by the subduction of a thinned continentalmargin) was not precisely constrained, but at leastHP conditions (depths >40 km) were reached(e.g. Kr€oner et al., 2012; Rojas-Agramonte et al.,2013). Furthermore, a suture marking the collisionbetween the North Tianshan and the Anrakhai mi-crocontinents running ~50 km in E–W directionoccurs to the north of the Aktyuz Complex (Bakirov& Maksumova, 2001; Ghes’, 2008; Kr€oner et al.,2013; Rojas-Agramonte et al., 2013; Klemd et al.,2014b).This model involving a ‘slab break-off’ to explain

the granitoid intrusions was proposed for the forma-tion of the Anrakhai and Aktyuz Complexes, butexcluding the Makbal Complex (cf. Kr€oner et al.,2012). Even though metamorphic ages of the Aktyuzand Makbal metamorphic complexes are very similar(Aktyuz: 474 � 2 Ma, Rojas-Agramonte et al., 2013and 462 � 7 Ma, Klemd et al., 2014b; Makbal: HPamphibolite: 470.1 � 2.5 Ma, Rojas-Agramonteet al., 2013 and UHP Grt-Cld-Tlc schist: 475 �4 Ma, this study) and although it is tempting torelate their formation to the same geodynamic envi-ronment, their position in the North Tianshanmicrocontinent collage rather points to two differentsuture zones: The Aktyuz metamorphic complexrelated to the subduction of the Djalair-Naiman basinis situated at the northeastern margin of the NorthTianshan (Fig. 1, Djalair-Naiman Suture Zone; e.g.Kr€oner et al., 2012). On the contrary, the MakbalComplex is related to subduction under the south-western margin of the North Tianshan and is situatednear the southwestern margin of the North Tianshan.Given that the affiliation of the microcontinent south-west of the Makbal area (Karatau-Talas terraneaccording to Kr€oner et al., 2013 or Ishim-Middle-Tianshan microcontinent after e.g. Windley et al.,2007) to either the Middle Tianshan (e.g. Windleyet al., 2007; Qian et al., 2009; Biske & Seltmann,2010; Alexeiev et al., 2011; Seltmann et al., 2011) orthe North Tianshan (e.g. Gao et al., 2009; Orozbaevet al., 2010; Konopelko et al., 2012; Kr€oner et al.,2014) is not clear yet (e.g. Fig. 1 and discussion inKr€oner et al., 2013), the ocean associated with the



formation of the Makbal Complex is either the Ky-rgyz-Terskey Ocean (Nikolaev Line suture zone) or asmaller oceanic (back-arc) basin between the micro-continent southwest of Makbal and the North Tian-shan.

The Makbal Complex forms an antiformal struc-ture, a so-called ‘brachi-antiform’ (cf. Konopelkoet al., 2012; Degtyarev et al., 2013). This structuralantiform or ‘dome’ consists of (U)HP metasedimenta-ry rocks with (U)HP mafic lenses and boudins of oce-anic origin (e.g. Meyer et al., 2013) with garnetgrowth ages of 475 � 4 Ma (this study). This (U)HPdome is surrounded by lower grade metamorphicrocks: Mafic rocks of continental origin (mafic dykesintruded in marginal continental crust) sampled~13 km south of the central Makbal location(e.g. Rojas-Agramonte et al., 2013: samples KG79and 79a) record a significantly different P–T evolu-tion and lower pressure conditions than the (U)HProcks from the centre, thereby defining a pressuregap between the (U)HP ‘dome’ and surroundingrocks at Makbal. The existence of thrusts or normalfaults between these two sampled occurrences wasverified by recent mapping (thrusts and folds, nappestructures, Degtyarev et al., 2013) and our fieldworkat the Makbal Complex.

In addition to differences in maximum pressureconditions, the determined geothermal gradient ofthe Makbal subduction zone shows large variations.Recent results of Rojas-Agramonte et al. (2013) sug-gested a ‘hot’ thermal gradient for the HP continen-tally derived rocks sampled ~13 km south of thecentral Makbal locality. Other studies on mafic (oce-anic) rocks from the central Makbal locality (e.g.Meyer et al., 2013) suggested cold subduction as indi-cated by the preservation of significant amounts ofpeak metamorphic hydrous minerals like white micaand/or glaucophane in the eclogite facies metabasites.This is also especially true for the mafic Grt-Cld-Tlcschist (this study) that contains a very high amountof stable hydrous UHP minerals (talc, chloritoid,phengite). Thus, all mafic (oceanic) rocks from thecentral Makbal Complex show evidence for a coldsubduction zone, and that a significant amount ofhydrous minerals was stable under HP–UHP condi-tions. Such differences concerning not only therespective geothermal gradient but also the P–Tpaths derived from thermodynamic modelling (cf.figs 11 & 14 in Meyer et al., 2013 and fig. 9b in Ro-jas-Agramonte et al., 2013) may be attributed tochanges in the subduction system (oceanic v. conti-nental subduction and resulting differences in slabdip and thus different geothermal gradients, e.g.Carry et al., 2011; Klemd et al., 2011).

The exhumation of (U)HP oceanic crust and HPmafic continental rocks and especially the occurrenceof a pressure-gap is best explained by an ‘UHPdome’ exhumation model of continental crust (War-ren et al., 2008a,b; Beaumont et al., 2009). These

authors suggested a geodynamical model for crustalburial of the thinned continental (former passive)margin during the initial c. 10 Ma of continental col-lision and exhumation (thus, the continental-typesubduction and exhumation of Guillot et al., 2009 orgenerally alpine-type subduction) in a subductionchannel below an accretionary wedge. An UHPdome is thought to be exhumed rapidly in a pulse-like buoyant up-channel flow from the deep (up to>100 km) subduction channel as soon as the exhu-mation number (E, involving not only buoyancy, butalso viscosity, channel thickness and other factors,e.g. Raimbourg et al., 2007; Warren et al., 2008a,b)is >1 (Beaumont et al., 2009). The results of numeri-cal modelling of Beaumont et al. (2009) are summa-rized in the following: After limited slab retreat dueto collision (i.e. onset of subducting the continentalmargin following oceanic subduction), the thinnedcontinental margin is subducted beneath the accre-tionary wedge. While UHP material is accumulatedin the lower part of the subduction channel, a nappeof subducted material (non-UHP conditions) decou-ples from the downgoing slab in the upper part ofthe subduction channel. As soon as the leading edgeof the strong continental interior reaches UHP condi-tions, E increases rapidly and the accumulated UHPmaterial converts to a rising UHP plume movingupwards rapidly. During the up-channel-flow, thelower grade nappe is underthrust by the UHP domeat lower crustal depth and both move further up.Once the buoyant dome has reached the base of theaccretionary wedge, the latter becomes destabilizedresulting in lateral extension of the wedge accom-plished by syn-exhumation normal faulting. Afterexhumation, the UHP dome is overlain and sur-rounded by thrust-stacks of lower grade subductedrocks and material of the accretionary wedge. Dif-ferent units of the thrust-stacks are separated bynormal faults and the suture zone is located close tothe upper continental plate. Thus, this model fits tothe tectonic situation at Makbal where the transitionfrom oceanic subduction to subduction of thethinned continental margin (including mafic dykes)is indicated from the exhumed (U)HP rocks. Theexhumation mechanism of oceanic crust in a sub-duction channel due to forced channel flow intro-duced above for formation of the central Makbal(U)HP tectonic m�elange is still valid, because the‘dome’ exhumation model accounts for the exhuma-tion during early continental collision. However, thefast uplift of the continental margin may assist and/or accelerate the exhumation of discontinuous frag-ments of oceanic crust, which are already juxtaposedin a presumably sedimentary subduction channel.Thus, the central Makbal Complex may representthe UHP dome, and the lower pressure samplesouth of Makbal (Rojas-Agramonte et al., 2013) isthought to represent a section of the HP-nappe (cf.Degtyarev et al., 2013).



CONCLUSIONS

At Makbal, strongly metasomatized metabasic rocks(either strongly hydrothermally altered oceanic crust orarc-related volcaniclastics) now preserved as a Grt-Cld-Tlc schist were subducted to UHP conditions reachingmaximum pressure conditions of ~2.85 GPa and maxi-mum temperatures of ~580 °C with a prograde to peakmetamorphic Sm-Nd garnet age of 475 � 4 Ma.

A strong coupling between garnet growth andwater release during prograde metamorphism can bededuced from thermodynamic modelling of the Grt-Cld-Tlc schist, confirming the observation of Baxter& Caddick (2013) that garnet growth may be appliedas a proxy for dehydration reactions in a subductingslab.

The existence of a (U)HP tectonic m�elange at thecentral Makbal Complex with relicts of oceanic crust(and arc-related volcaniclastic material) embedded involuminous continental (clastic) metasedimentaryrocks indicates initial exhumation and juxtapositionof incoherent pieces of oceanic crust derived fromdifferent depths in the subduction channel duringcontinuing subduction. Subsequent subduction ofmarginal continental crust is indicated by the occur-rence of HP continental mafic rocks with similarmetamorphic ages, but a different P–T path and geo-thermal gradient ~13 km to the south (Rojas-Agramonte et al., 2013).

The exhumation of deeply subducted (U)HP rocksat Makbal either is due to slab-rollback (e.g. Hacker& Gerya, 2013) or channel flow (e.g. Hacker &Gerya, 2013; Hacker et al., 2013). The latter mecha-nism explains both the occurrence of domains withdifferent subduction histories and the presence of astructural ‘dome’ cored by UHP rocks surroundedby lower pressure nappes and associated with volu-minous metasedimentary rocks accumulated in theaccretionary wedge, which acts as host for the domedriven by the emplacement of an exhuming nappestack (Warren et al., 2008a,b; Beaumont et al.,2009).

ACKNOWLEDGEMENTS

This research was supported by the Deutsche Fors-chungsgemeinschaft (KL692/17-3). DK was sup-ported by research grants from Saint PetersburgState University 3.38.137.2014 and 3.39.139.2014.We thank U. Sch€ussler, H. Br€atz and M. Koch-M€uller for assistance during laboratory work. O.Vokueva helped in gaining the field-data and C. Be-ier and K. Haase are thanked for very helpful dis-cussions. Detailed and constructive reviews of A.Vitale Brovarone and A. Okay helped to clarify themanuscript and are highly appreciated. D. Robinsonis thanked for editorial handling and detailed com-ments. The research is a contribution to the IGCP-592 project.

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SUPPORTING INFORMATION

Additional Supporting Information may be found inthe online version of this article at the publisher’sweb site:Figure S1. PRIMA-normalized trace element data

of the metasedimentary samples from Makbal. Forcomparison (grey shaded area), other metasedimenta-ry rocks from Makbal and passive margin metasedi-mentary rocks compiled from McLennan et al.(1990) are shown.Figure S2. PRIMA-normalized trace element data

of volcanic glass analyses of the East Scotia Ridge(back-arc basin) from Fretzdorff et al. (2002). Thesesamples show that in back-arc basins, a negative Nb-Ta anomaly may be obvious (green lines), poorlydeveloped (blue lines) or even missing (grey lines).Table S1. Major and trace element data of other

mafic and metasedimentary rocks from Makbal anddata from McLennan et al. (1990) for sedimentaryrocks shown for comparison.Table S2. Mineral compositions used for avPT

(sample 10-16).

Received 26 February 2014; revision accepted 19 June 2014.



Subduction and exhumation mechanisms of ultra-high and high-pressure oceanic and continental crust...

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