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Invited Review Article
Forty years of TTG research
Jean-François Moyen a,c,d,⁎, Hervé Martin b,c,d
a Université Jean-Monnet, Université de Lyon, 23 rue du Docteur Michelon, 42023 Saint-Etienne, Franceb Clermont Université, Université Blaise Pascal, Laboratoire Magmas et Volcans, 5 rue Kessler, F-63038 Clermont-Ferrand Cedex, Francec CNRS, UMR 6524, LMV, F-63038 Clermont-Ferrand, Franced IRD, R 163, LMV, F-63038 Clermont-Ferrand, France
Article history:Received 31 October 2011Accepted 8 June 2012Available online 16 June 2012
Keywords:ArchaeanTTGPetrologyGeochemistry
TTGs (tonalite–trondhjemite–granodiorite) are one of the archetypical lithologies of Archaean cratons. Sincetheir original description in the 1970s, they have been the subject of many studies and discussions relating toArchaean geology. In this paper, we review the ideas, concepts and arguments brought forward in these40 years, and try to address some open questions — both old and new.The late 1960s and the 1970s mark the appearance of “grey gneisses” (TTG) in the scientific literature. Duringthis period, most work was focused on the identification and description of this suite, and the recognitionthat it is a typical Archaean lithology. TTGs were already recognised as generated by melting of maficrocks. This was corroborated during the next decade, when detailed geochemical TTG studies allowed us toconstrain their petrogenesis (melting of garnet-bearing metamafic rocks), and to conclude that they musthave been generated by Archaean geodynamic processes distinct from their modern counterparts. However,the geodynamic debate raged for the following 30 years, as many distinct tectonic scenarios can be imagined,all resulting in the melting of mafic rocks in the garnet stability field. The 1990s were dominated by experi-mental petrology work. A wealth of independent studies demonstrated that melting of amphibolites as wellas of mafic eclogites can give rise to TTG liquids; whether amphibolitic or eclogitic conditions are more likelyis still an ongoing debate. From 1990s onwards, one of the key questions became the comparison with mod-ern adakites. As originally defined these arc lavas are reasonably close equivalents to Archaean TTGs.Pending issues largely revolve around definitions, as the name TTG has now been applied to most Archaeanplutonic rocks, whether sodic or potassic, irrespective of their HREE contents. This leads to a large range ofpetrogenetic and tectonic scenarios; a fair number of which may well have operated concurrently, but are ap-plicable only to some of the rocks lumped together in the ever-broadening TTG “bin”.
Fig. 1. A “sea of granites”: Archaean geology in the late 1970s. On this geological map ofthe Baltic shield [after Salop and Scheinmann (1969) and Gaál et al.(1978) in Condie(1981)], Archaean terrains were subdivided into two lithologies: greenstone beltsand granito-gneissic basement. The latter was considered as a “sea” of granites andgneisses, without further distinction between these felsic lithologies.
1. Introduction
The first Archaean units investigated were greenstone belts, as theyinclude contrasting lithologies, such as komatiites or banded iron forma-tions, which are potentially ore-bearing and of economic interest. Themore homogeneous and economically less promising granitoidsremained almost unknown for a long period. Glikson (1979) statedthat “The largely unknown nature of the granitic terrains is manifested bytheir presentation as undivided pink areas on regional geological mapsand their designation in such terms as ‘sea of granite’ (Yilgarn, WesternAustralia) or ‘Peninsular Gneiss’ (Southern India)” (Fig. 1). Only at theend of the 1960s and beginning of the 1970s did geologists begin toshow interest in the “sea of grey gneiss” (Anhausser et al., 1969; Blissand Stidolph, 1969; Glikson and Sheraton, 1972; Heimlich, 1969;Heimlich and Banks, 1968; Hunter, 1970; Lund, 1956; McGregor, 1973;Sheraton, 1970; Viljoen andViljoen, 1969). The development of geochro-nology, initially by K–Ar and Rb–Sr methods (Black et al., 1971; Goldichet al., 1970; Hanson et al., 1971; Heimlich and Banks, 1968; Kouvo andTilton, 1966; Moorbath, 1975; Moorbath et al., 1972) led to the first ab-solute ages of the different components of Archaean granitoid base-ment. The first studies, for technical reasons such as samplefreshness, or quality of the outcrops, focussed on late granites: typi-cally potassic granites cutting across the surrounding gneisses. How-ever, this was a major milestone, fixing reliable temporal markers forall Precambrian terrains — that is for about 88% of Earth history. Atthe same time, the petrographic and geochemical characterisationof the grey gneiss suites had been undertaken (Arth and Hanson,1975; Barker et al., 1979; Bridgwater and Collerson, 1976; Condieand Hunter, 1976; Hanson and Goldich, 1972; Hunter et al., 1978;O'Nions and Pankhurst, 1978; Tarney et al., 1979; Weaver andTarney, 1980, among the pioneers), revealing the close associationof three (sodic) plutonic types: tonalites, trondhjemites and granodio-rites. This characterisation led to use of the acronym TTG, first used in apublication by Jahn et al. (1981).
a
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Since the 1980s, TTGs have been widely studied, with two mainfocuses: (1) understanding the tectonic regime of the Early Earthand (2) constraining the processes of differentiation of the continen-tal crust. The aim of this paper is to review the history of conceptsand questions relating to TTGs, broadly in chronological order.Many of these questions remain incompletely answered, and eachsection will therefore contain several questions and discussionpoints. At the end, we will outline a few further questions recentlyidentified, that remain to be discussed.
Qb
2. 1970s: characterisation of TTG suites
The pioneer studies on TTGs, in the late 1960s and 1970s, focussedon the description of these rocks, and their comparison with othercommon granitoids.
PA
a
T
b
cA
Fig. 3. Modal composition of TTGs. a) In thin section, typical TTG consist of an assem-blage of quartz+oligoclase+biotite, this later mineral carrying the metamorphic foli-ation (2.9 Ga-old Kivijärvi TTG, Finland; photo H. Martin; width of photo: ca. 5 mm);b) Q–A–P modal classification of Streicksen (1975) for the 2.9 Ga-old Kivijärvi TTG,showing their mostly tonalitic and granodioritic composition (grey field=tonalite+granodiorite). The points also plot along the K-poor calc-alkaline differentiationtrend of Lameyre and Bowden (1982), (T=tholeiitic suite, A=alkaline suite; thecalc-alkaline suites are subdivided into: a=K-poor; b=intermediate; c=K-rich).
2.1. Field data and petrographic characteristics
TTGs are the main components of Archaean terrains. They mostlycrop out as diversely deformed plutonic rocks, ranging from pristineor nearly pristine plutons (Fig. 2a); to orthogneisses (Fig. 2b); to com-ponents of complex, banded gneiss complexes, often migmatitic andcrosscut by mafic and/or granitic dykes (“grey gneisses”: Fig. 2c, d).TTGs have a typical quartz+oligoclase+biotite mineral assemblage(Fig. 3a). K-feldspar (microcline) is very scarce, whilst the less differen-tiated rocks of the suite contain hornblende. Themain accessory phasesare allanite, pistacite, apatite, zircon, titanite and titanomagnetite; thepresence of magmatic epidote is very significant, as it implies thatcrystallisation started in the mid- or lower crust (Schmidt and Poli,2004; Schmidt and Thompson, 1996). The typical modal composition(Fig. 3b) is tonalitic (Streckeisen, 1975), and defines a K-poor calc-alkaline differentiation trend (Lameyre and Bowden, 1982).
a b
c d
Fig. 2. Field examples of TTGs: from plutons to banded gneisses. a) 3.45 Ga-old homogeneous grey gneisses (TTG) of Stolzburg (Barberton area, S. Africa); b) 2.7 Ga-old homoge-neous grey gneisses (TTG) of Naavala (Kainuu, Finland); c) 3.1 Ga-old banded heterogeneous grey gneisses of Sand River (Limpopo, S. Africa); d) 3.64 Ga-old banded Ancient GneissComplex (Swaziland).
Table 1Proposed definition of TTGs and related rock types (cf. Table 1). The grey boxes correspond to criteria matching the (proposed) definition of TTG s.s. All elements of this definition are equally important, and no rock should be called TTGunless it matches all features (and not only a handful, such as La/Yb or Sr/Y ratios). This definition deliberately excludes (i) non-plutonic portions of grey gneisses; (ii) potassic plutonic rocks; (ii) low-Al2O3 and/or high-HREE sodic plutonicrocks (low pressure “TTGs” in Moyen, 2011a).
low HREE med. HREE high HREE
Trondhjemites common Tonalites common
Major minerals - felsic Quartz + Oligoclase + K-feldspar
Magmatic epidote commonHornblende commonly
abundant
Accessory minerals
68 < SiO2 < 72 % 65 < SiO2 < 70%
Fe2O3*+ MgO + MnO + TiO2 <
4%
K2O ca. 3%
3 < Na2O < 5 % Na2O > 5%
0,8 < K2O/Na2O < 1
Aluminium > 14 % at 72% SiO2 > 16 % at 70% SiO2 > 15% at 70% SiO2 > 14 % at 70% SiO2
Fig. 4. Major element features of TTGs. a) Normative An–Ab–Or triangle (Barker, 1979;O'Connor, 1965) showing that the composition of Archaean TTG (blue points) differsfrom that of the modern continental crust (yellow field); b) K–Na–Ca triangle contrast-ing the evolution of TTGs and modern calc-alkaline magmas (brown CA trend). CAtrend is characterised by K-enrichment during differentiation, whereas TTGs show noclear trend and remain Na-rich and K-poor.
(YbN)
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00 4 8 12 16 20
LaN
CeLa Nd SmEuGdTbDy Er YbLu
1
10
100
RO
CK
/ C
HO
ND
RIT
ES
Archaean TTG
1CeLa Nd SmEuGdTbDy Er Yb Lu
Post 2.5 Ga granitoids
2
RbBaTh KU Nb LaTa Ce Sr NdZr EuGdSmHf Ti Dy Y Er Yb V Cr Ni
0.1
0.01
1
10
100
RO
CK
/ P
RIM
ITIV
E M
AN
TL
E
a
b
Fig. 5. Trace element features of TTGs. a) (La/Yb)N vs. YbN plot for Archaean TTG (bluepoints) and modern continental crust (yellow points). A clear compositional changeappears at the Archaean–Proterozoic boundary (~2.5 Ga). Archaean TTG (Inset 1)have low-HREE contents (0.3bYbNb8.5) associated with strongly fractionated patterns(high (La/Yb)N), while the post-2.5 Ga continental crust has high-HREE contents(4.5bYbNb20) and moderately fractionated REE patterns ((La/Yb)N≤20) (Inset 2).Normalisation values after Masuda et al. (1973). b) Primitive mantle normalised(McDonough et al., 1992) trace element patterns for both TTG (blue pattern) and mod-ern continental crust (brown).
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2.2. Geochemical characteristics
TTGs are silica-rich (SiO2>64 wt.%, but commonly ~70 wt.% orgreater) with high Na2O contents (3.0–7.0 wt.% Na2O) and correlatedlow K2O/Na2O (b0.5). They are poor in ferromagnesian elements(Fe2O3*+MgO+MnO+TiO2≤5 wt.%, Fe2O3*=total Fe expressed asFe2O3), with an average Mg# of 0.43 and average Ni and Cr contentsof 18 and 40 ppm respectively (Table 1). In the normative classificationdiagram for granitoids (Fig. 4a) (O'Connor, 1965), the plutonic compo-nent of Archaean grey gneisses all plot in the trondhjemite, tonalite andgranodiorite domains, thus accounting for the classically used TTG acro-nym (Jahn et al., 1981). This is in sharp contrastwithmodern granitoids,which are commonly richer in K2O (granodiorites to granites). Similar-ly, as shown in the cationic K–Na–Ca triangle, granitoids from themodern continental crust evolve through K-enrichment during differ-entiation, while Archaean TTGs do not show a clear differentiationtrend, but plot close to the Na apex (Fig. 4b). Based on Al2O3 content,Barker and Arth (1976) subdivided trondhjemites into high- and low-Al2O3 groups;most Archaean TTG having Al2O3>15% at SiO2=70%, be-long to the high‐Al2O3 group. They evolve from metaluminous toslightly peraluminous (A/CNK=1; normative corundumb1%),with A/CNK increasing towards silica-richer compositions.
TTGs exhibit characteristic trace element signatures, with high con-tents in light Rare Earth Element (LREE) (Laaverage=31.4 ppm) but verylow heavy Rare Earth Element (HREE) contents (Ybaverage=0.64 ppm),resulting in high La/Yb ratios ((La/Yb)average=49, corresponding to(La/Yb)N=32.4; chondritic REE normalisation values are from
Masuda et al. (1973) divided by 1.2). They also lack significant Euand Sr anomalies (Fig. 5a, Martin, 1986) but do show negativeNb–Ta and Ti anomalies (Fig. 5b); Nb/Ta ((Nb/Ta)average=8.7)and Lu/Hf ((Lu/Hf)average=0.033) ratios are low while Zr/Sm ratiosare high ((Zr/Sm)average=44.3). The lack of both Eu and Sr anoma-lies associated with the low HREE content are interpreted asreflecting the presence of garnet and amphibole as well as thelack of plagioclase, either as residual or fractionating phases; suchthat, most often, a high-pressure origin has become implicit inthe term ‘Archaean TTG’ (Champion and Smithies, 2003). In this re-spect too, this trace element signature makes TTG different frommodern continental crust, which has both higher HREE contents(Ybaverage=3.26 ppm), and negative Eu, Nb–Ta, Sr and Ti anoma-lies (Fig. 5b).
2.3. Definition issues
Although early definitions of TTGswere relatively unambiguous, thename has been usedmuchmore loosely in the literature. Tables 1 and 2summarise the nature and composition of the various rock types thathave been called “TTG”; in Table 1, the grey boxes highlight the featuresmatching the definition of Sections 2.1 and 2.2. TTGs often occur as acomponent of the ubiquitous “grey gneiss complexes” that form the“background” lithology of most Archaean cratons. Although the term“grey gneisses” is commonly taken as being synonymous to TTG, theformer contains a range of tectonically transposed components
Table 2Average composition of (1) Archaean grey gneisses, unsorted; (2) plutonic, plagioclase-dominated component of the grey gneisses; (3) potassic and (4) sodic (=TTG s.l.) plutoniccomponents, the sodic TTGs being themselves subdivided into three groups. Authors' database (Martin and Moyen, 2002; Moyen, 2011a). Mg#=cationic ratio Mg/Mg+Fe; A/CNK=cationic ratio Al/2Ca+Na+K; (La/Yb)N=chondrite-normalised La/Yb; Eu=Eu� ¼ EuN
ffiffiffiffiffiffiffiffiffiffiffiffiffiffi
SmN :GdNp , the N subscript indicates values normalised to chondrites (Masuda et al., 1973).
Major elements are given in weight% and trace elements in ppm. Avg=average (bold); std=standard deviation.
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(Fig. 2c, d), not all of them being metagranitoids. Most outcrops areheterogeneous and not only contain TTG, but also K-rich granites,leucosomes, restites, and even amphibolites or metapelites — allencompassed under the global description of “grey gneisses”. In ad-dition, the plutonic components of grey gneisses actually include awide range of rocks (Moyen, 2011a) that comprise both sodic andpotassic igneous rocks — of which only the sodic components areTTGs, by any reasonable definition. Similar potassic plutonic rocksare also emplaced as individual plutons, sometimes referred to as“enriched TTGs” or “transitional TTGs”. Finally, even the sodic plu-tonic (or metaplutonic) rocks include both low- and high-Al2O3
compositions, typically correlated with more or less fractionatedREE patterns. Strictly speaking, only the high-Al2O3 rocks with frac-tionated REE patterns should be referred to as TTGs, but the general
usage unfortunately became much looser (e.g. Feng and Kerrich,1992; Whalen et al., 2002; Willbold et al., 2010).
Unless specified otherwise, wewill restrict the use of the term “TTG”to the high-Al2O3 sodic plutonic or metaplutonic igneous rocks withfractionated REE patterns and low HREE contents (i.e. low andmediumHREE sodic in Table 2). “TTG s.l.”will be used in a broader sense to referto both low- and high-Al2O3 sodic (meta-)plutonic rocks.
3. 1980s: petrogenetic modelling based on geochemistry
At the end of the 70s and beginning of the 80s, several petrogenet-ic scenarios were proposed to account for the geochemical character-istics of Archaean TTG.
200
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100
50
00 10 20 30
500
400
300
200
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00 10 20 30 40 50
a
b
PaleogenePliocenePleistocene
Limit of fig. 6a
Defant & Drummond’s field of adakites
Defant & Drummond’sfield of adakites
Mc Pherson et al.field of adakites
High-P frac. cryst.of arc basalt
Arc basalt melting
Slab melting50
70
80
50
80
50
Sr/Y
Y
Sr/Y
Y
Fig. 6. Fractionation of commonmafic parents to give high Sr/Y melts (Macpherson et al.,2006). a) Variousmodels for the origin of Tertiary adakites inMindanao (Philippines). Slabmelting, arc basalt melting and fractionation of a basaltic parent were all tested (curves;ticks and numbers correspond to the percentage of liquid remaining) and are all viablemodels. An isotopic mismatch between the slab and the adakites, as well as a correlationbetween SiO2 and Sr/Y, are taken as further evidence favouring fractionation over a slabmelting model (MacPherson et al., 2006). Note, however, the difference between theadakitic field as used in this figure; and the original field (Defant and Drummond,1990), drawn in panel b. Also note that Mindanao adakites do not reach the very highSr/Y values observed in many other modern adakites (grey field in panel b).
318 J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336
3.1. “Historical” models for TTG genesis
The four main models proposed to account for the origin of TTGswere:
1) Fractional crystallisation of awet basalticmagma leaving a cumulatemade up of hornblende±plagioclase±biotite (Arth, 1979; Arthet al., 1978; Barker, 1979; Maaløe, 1982; Smith et al., 1983). Howev-er, this model requires more than 75% fractional crystallisation togenerate trondhjemitic liquids from a basaltic source and shouldtherefore produce enormous amounts of cumulate (Arth et al.,1978; Martin et al., 2005; Meijer, 1983; Spulber and Rutherford,1983). In addition, fractionation would generate continuous trendsfrom the parental (basalt) to the differentiated liquid, and suchtrends are not observed in TTG suites. Finally, a fractionationmodel requires concomitant basaltic and TTG magmatism, both intime and space. Although Archaean cratons do include significantamounts of mafic rocks (in the form of lavas from greenstonebelts), detailed geochronology demonstrates that the TTGs and thebasalts are seldom coeval. In fact, TTG emplacement is concomitantwith breaks in themafic activity in the nearby greenstone belts, dur-ing which the erupted lavas are rather felsic and similar to the TTGplutons (Benn and Moyen, 2008; Champion and Smithies, 2007;Moyen et al., 2007; Smithies et al., 2007).This model was revived in the 2000s to explain the evolution of acompositionally similar group of rocks called adakites (seeSection 5 below). Amphibole (Arth et al., 1978; Davidson et al.,2007; Kleinhanns et al., 2003) or, more significantly, garnet(Alonso-Perez et al., 2009; Macpherson et al., 2006; Prouteau andScaillet, 2003) fractional crystallisation was proposed to play a rolein shaping the geochemistry of some arc suites (Richards andKerrich, 2007). While amphibole has only a limited potential to cre-ate the typical HREE depletion observed in adakites/TTGs (Davidsonet al., 2007), garnet fractionation is a muchmore significant processthat has been demonstrated both geochemically (Macpherson et al.,2006) and experimentally (Alonso-Perez et al., 2009) to be capableof generating HREE-depleted intermediate magmas from an ordi-nary andesitic parent at pressures of ca. 10 kbar (Fig. 6). While thisprocess is undoubtedly able to generate some high Sr/Y or La/Ybmagmas (Moyen, 2009), its applicability to TTGs is somewhat debat-able. Firstly TTGs aswell as adakites do not show long differentiationtrends and the “less differentiated” (basalt to basaltic andesite) end-members are typically missing. Secondly, an important feature ofboth adakites and TTGs is that their K/Na ratio remains low through-out differentiation, and does not correlate with differentiation indi-cators such as SiO2. On the contrary, whenever fractionation ofamphibole or garnet is suspected (Macpherson et al., 2006;Richards and Kerrich, 2007) or experimentally observed (Alonso-Perez et al., 2009), the melts invariably become more potassic(also see Fig. 14 below) in the course of differentiation. Finally, ingeneral one may dispute the similarity between plutonic and volca-nic rocks, the latter being typically dryermagmas, withmore poten-tial to ascend, erupt, and undergo differentiation (Clemens andDroop, 1998).
2) Direct melting of the mantle, possibly metasomatised by fluids(Moorbath, 1975; Peterman and Barker, 1976). In this case verylow degrees of melting (b5%) are required to generate felsicmagmas. However, even at such low degrees of melting, theoreticalcalculations based on REE concentrations demonstrate that it is im-possible to account for the Yb depletion and the high La/Yb ratiostypical of Archaean TTG (Fig. 7 and Jahn et al., 1981, 1984; Martin,1987; Martin et al., 1983). Furthermore, a quartz-normativemagma, such as TTG, is unlikely to be in equilibriumwith ultramaficmantle.
3) Partial melting of Archaean greywackes (Arth and Hanson,1975). Although this model of recycling explains the REE and
some trace-element characteristics of Archaean TTG, it cannotaccount for major- and other trace-element behaviour (Martin,1994). The sodic nature of TTGs in particular is at odds withthis model.
4) Partial melting of a hydrated basalt metamorphosed at high pres-sure and transformed into eclogite or garnet-bearing amphibolite(Arculus and Ruff, 1990; Arth and Hanson, 1975; Barker, 1979;Barker and Arth, 1976; Compton, 1978; Condie, 1981, 1986;Condie and Lo, 1971; Ellam and Hawkesworth, 1988; Glikson,1979; Gower et al., 1983; Hanson and Goldich, 1972; Hunteret al., 1978; Jahn et al., 1981, 1984; Martin, 1986, 1987, 1988;Martin et al., 1983; Nédelec et al., 1990; Sheraton and Black,1983; Tarney et al., 1982). This model became generally acceptedin the late 80s, with the further requirement that garnet must bestable in the residue in order to account for the HREE depletionof TTG, putting strong constrains on the depth of melting. Fig. 7shows the key (geochemical) elements of this model.
3.2. The reference model: melting of meta-mafic rocks
By the end of the 1980s, a “reference model” accounting for TTGgenesis had progressively evolved and matured into a 3-stage mech-anism (Fig. 8):
(YbN)8 12 164
E
G10
G0
G25
50
10101010
25
252525
101010
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AT
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La Ce Nd SmEu Gd DyTb Er Yb Lu
RO
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/ C
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ES
50
5
La Ce Nd SmEu Gd DyTb Er Yb Lu
Source = G 0
Archaean tholeiite
TTG
F = 10%
F = 50%
Calculated magmacomposition
Source = G 25
Archaean tholeiite
TTG
F = 10%
F = 50%
Calculated magmacomposition
2 3
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0 8 12 16
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La Ce Nd SmEu Gd DyTb Er Yb Lu
RO
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5 Source = PmPrimitive Mantle
TTG
F = 1%
F = 25%
Calculated magmacomposition
10
25
L1
L3L2
10
10
25 25 3020 2030
PmMm
1
a b
Fig. 7. Geochemical tests of TTG petrogenetic models. (a) (La/Yb)N vs. YbN diagram summarising different models for mantle melting (Martin, 1986, 1987). Both primitive (Pm) and met-asomatised (Mm)mantle compositions correspond tomelt in equilibriumwith a lherzolite residue containing 10% (L1) and 5% (L2) garnet or 5% spinel (L3). Dotted lines are olivine fractionalcrystallisation trends at depth. Numbers indicate the degree of melting or of fractional crystallisation. Blue field=Archaean TTG. Inset (1) represents themodelled REE patterns formelting of achondriticmantle source (Martin, 1987). Grey pattern=Primitivemantle (Pm); blue pattern=average TTG; pale yellow field=domain of REE patterns for liquids generated by 1 to 25%melt-ing of the source. (b) (La/Yb)N vs. YbN diagram summarising the differentmodels for basalt melting. Source is an average Archaean tholeiite (AT) transformed into garnet-free amphibolite (G0;inset 1); 10% (G10) and 25% (G25; inset 2) garnet bearing amphibolite aswell as into eclogite (E). Blue field=Archaean TTG. Insets represents the REE patternsmodelled for G0 (inset (2)) andG25 (inset (3)) (Martin, 1987). Grey pattern=Archaean tholeiite; blue pattern=TTG; pale blue and redfield=domain of REE patterns for liquids generated by 10 to 50%melting of the source.
40
30Source
Rutile-eclogite
319J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336
1) partial melting of the mantle to generate a basalt;2) melting of this basalt, metamorphosed to garnet bearing amphib-
olite or eclogite, to give rise to TTG parental magma;3) limited low pressure (intacrustal) fractional crystallisation to pro-
duce the differentiated TTG suites.
The third step – fractional crystallisation – may or may not haveaffected all suites.Where it has occurred, it was in the form of extrac-tion of hornblende±plagioclase from the parental TTGmagma. In allcases, the degree of fractionation was lower than 25% (Martin, 1987;Moyen et al., 2007); however, it is important in shaping the overallgeochemical trend of a given pluton or suite.
3.3. Melting of meta-basalts: at what depth?
Later refinement of the general model persisted for the next two de-cades. For example, while it is widely accepted that TTGs were generat-ed by melting of hydrous meta-basalt, the depth of melting remainedhotly debated, with either a garnet–amphibolite or an eclogite residueappearing to be realisticmodels. Foley et al. (2002) investigated the dis-tribution of HFSE in TTG liquids and concluded that the low Nb/Ta and
Stage 1
Stage 2
Stage 3
TONALITIC MAGMAResidue
Hbl+Grt+Cpx+Ilm±Pl
CumulateHbl+Ilm±Pl
THOLEIITE
MANTLE
FC
PM
PM
T.T.G. SUITE
Fig. 8. Schematic diagram summarising the succession of the different mechanisms im-plied in Archaean TTG genesis (after Martin, 1993). Hbl=hornblende; Grt=garnet;Pl=plagioclase; Cpx=clinopyroxene; Ilm=ilmenite; PM=partialmelting; FC=fractionalcrystallisation.
high Zr/Sm ratios in Archaean TTG preclude an eclogitic residue butare consistent with a garnet bearing amphibolite, assuming the sourceof TTGs had a chondritic Nb/Ta. Based on the same data, but with alow Nb/Ta basaltic source (not uncommon in the Archaean record),Rapp et al. (2003) arrived at the opposite conclusion (Fig. 9). We shallreturn to this question, after reviewing the results of experimental pe-trology studies.
4. The 1990s: experimental petrology confirms that melting ofmafic rocks in the garnet stability field gives rise to TTG melts
Following the consensus on the origin of TTGs by partial melting ofmeta-basalts, a large number of experiments were conducted fromthe 1990s onwards to investigate the feasibility of this scenario. A rela-tively recent paper (Moyen and Stevens, 2006) recorded in excess of
1 5 10 50 100 500 1000
0
10
20
Nb
/ T
a
Zr / Sm
Source AB-1(Rapp et al., 2003)
(Foley et al., 2002)
Eclogite
Amphibolite
1711717
252525
Fig. 9. Nb/Ta vs. Zr/Sm diagram comparing TTG composition (blue circles) with mod-elled composition. The red fields are from Foley et al. (2002), for melting of rutile bear-ing eclogite, rutile free eclogite and amphibolites; their source is chondritic (red star).These authors concluded that the Archaean TTG residue cannot be eclogitic. In contrast,based on a similar approach, but with a typical Archaean basalt composition (AB-1,green star), Rapp et al. (2003) obtained low Nb/Ta and high Zr/Sm (green field) byeclogite melting.
320 J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336
300 experiments in some 15 studies (Beard and Lofgren, 1991; Patiño-Douce and Beard, 1995; Rapp and Watson, 1995; Rapp et al., 1991;Rushmer, 1991; Sen and Dunn, 1994; Sisson et al., 2005; Skjerlie andPatiño-Douce, 1995, 2002; Springer and Seck, 1997; Winther, 1996;Winther and Newton, 1991; Wolf and Wyllie, 1994; Yearron, 2003;Zamora, 2000). All these studies clarified the details of TTG genesis,and essentially confirmed the globally accepted model.
4.1. A review of experiments, starting materials and water availability
The starting materials used in the experiments are of broadly ba-saltic composition, ranging from basalts to basaltic andesites and gen-erally belonging to the tholeiitic series, or being close to the calc-alkaline/tholeiitic boundary. Nevertheless, there are fairly significantdissimilarities between the different starting materials, in terms ofmodal composition, bulk rock chemistry and mineral chemistry.Amphibole/plagioclase ratios vary from0.18 to>4; quartz abundancesin experimental sources ranged from zero up to 24%. Other minerals(chlorite, epidote, Ti-oxides, titanite) were described in some startingmaterials. Accordingly, SiO2 content of the sources ranged from ca. 47
a
700 800 900 1000 1100 1200
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Tholeiitic basalt
P (
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)
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Trondhe
Granod
Granite
c
Source and water availQuartz-present, water Quartz-present, “dry”
Quartz-absent, water sQuartz-absent, “dry”
CaNa
K
Fig. 10. Main compositional features of experimental melts of amphibolites. a) In a K–Na–Cquartz in the source and the water availability, fluid-present or ‐absent) match the compositowards more calcic and alkali-poorer compositions (arrows). b) Focussing on one single setwith high pressure melts being more sodic, and low-pressure melts more calcic (Data from Zin the P–T space (Moyen and Stevens, 2006) for two groups of sources, tholeiitic (c, left) anexperiments. Three main controls can be observed on these diagrams: the role of the sourcature (high F, “hot” melts are more tonalitic) and the role of pressure (high pressure melts
to ca. 60%; K2O from 0.1 to 1.8% and Na2O from 1 to 4.3%. Mg# spreadbetween 38 and 71.
Experiments have been conducted over a fairly large P–T range,from 1 kbar to 35 kbar and from 750 to 1200 °C. Studies focusing onthe genesis of TTGs/adakites were mostly conducted in the garnet sta-bility field, i.e. greater than ca. 10 kbar. Nevertheless, the P–T spacehas not been completely explored, as most studies concentrated ona positively sloped “band” from 700–900 °C at Pb10 kbar to 1000–1100 °C at P>25 kbar. This implies that the behaviour of the systemis not described in some “exotic” parts of P–T space, such as theeclogitic region. Only a handful of recent studies tried to addressthis question (e.g. Laurie and Stevens, 2012; Skjerlie and Patiño-Douce, 2002).
Finally, different situations relative to water availability have beenexplored (Moyen and Stevens, 2006; Vielzeuf and Schmidt, 2001):(1) Fluid-present melting, with water either as a free phase (addedto the starting material) or released during prograde heating bybreakdown of hydrous phases such as chlorite or epidote. Dependingon the total amount of water initially present, the melt may or maynot have the potential to dissolve all of it, such that the system canbecome fluid-absent after the first melting increments. (2) Fluid-
b
40
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Arc basalt
700 800 900 1000 1100 1200
P (
kbar
)
T (°C)
jmite
iorite
SolidusPlag outAmp outGt in
CaNa
K
> 20 kbar
10 - 20 kbar
< 10 kbar
Pressureabilitysat.
at.
d
a triangle, most experimental melts (colour-coded based on the presence or absence oftion of TTGs (light blue field in the background). Increasing F values result in evolutionof experiments (with the same source and H2O contents), a pressure effect can be seenamora, 2000; plotted in Moyen, 2011a). c) and d) are “maps” of the melts compositiond arc-like (d, right). The coloured fields are modelled and the dots correspond to actuale (enriched source give more granitic/granodioritic compositions), the role of temper-are more trondhjemitic).
321J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336
absent melting, with all water in the system accommodated in hy-drous minerals (generally amphibole), whose destabilisation triggersmelting (“dehydration melting”). This is the most commonly investi-gated scenario, as the general consensus is that fluid-present meltingat >10 kbar (i.e., garnet stability field) is unlikely in nature; see how-ever Laurie and Stevens (2012) for an alternative view. (3) Totally drysystems, with no water at all even in hydrous minerals. This scenariomight be relevant to melting of a dehydrated eclogitic source, or to alower crustal granulite, and has hardly been investigated (Springerand Seck, 1997).
4.2. Major elements composition of melts
Despite the large range of sources and conditions, the results of allexperiments are fairly consistent. The position of the solidus wasfound to be controlled primarily by water availability (Moyen andStevens, 2006), while the melt fractions are controlled by the natureof the source and the amount of water in the system, with wet andquartz-rich sources being more fertile.
Since melting is essentially a eutectic process, the composition ofall experimental melts (regardless of the source composition) is rath-er similar. Broadly speaking, all melts (at reasonable temperatures)are leucocratic and sodic, and therefore very comparable to TTGs. Indetail, the following factors seem to affect melt compositions(Fig. 11):
(1) The degree of melting (F) is the most important parameter. As Fincreases, the melts become more mafic and calcic and K2O con-tents decrease. The behaviour of Na2O and Al2O3 depends on thenature of the source, these two elements being compatible inplagioclase-rich residues and incompatible otherwise. Therefore,with increasing F the liquids evolve from granitic or trondhjemitic(depending on the original K/Na ratio, Martin and Sigmarsson,2007), to granodioritic or tonalitic, to dioritic at very high(>50%) F values (Moyen and Stevens, 2006). Importantly, high Fcan result from high temperatures, but also from water orquartz-rich sources. Therefore melting of dry meta-basalts tends
Gar
net
0.00
0.20
0.40
40
30
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SolidusPlag outAmp outGt in
0
50
100
150
(La/
Yb
) N
40
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0
700 800 900 1000 1100
700 800 900 1000 1100
Tholeiitic basalt
P (
kbar
)
T (°C)
P (
kbar
)
T (°C)
Fig. 11.Mappingof garnet abundance (top) and La/Yb (bottom) in the PT space. The left column corrwith an arc-like trace elements pattern is used; this affects both the stability fields and abundance otherefore, of themelts). SeeMoyen and Stevens (2006) for further details. Garnet abundance is repr
to yield granitic, rather than trondhjemitic melts, as the lack ofwater inhibits melting (Prouteau, 1999).
(2) The composition of the source, and especially its K2O contents,is also significant. Indeed, in a systemwhere no potassic miner-al is present, K2O behaves as an incompatible trace element,such that its content in the melt is a function of both the sourcecomposition and the degree of melting (Shaw, 1970). More po-tassic sources thus yield potassic melts, granites and granodio-rites rather than trondhjemites and tonalites (Martin andSigmarsson, 2007; Sisson et al., 2005).
(3) Pressure exerts a marginal control. This effect is difficult toidentify, as it is hidden by the fact that increasing pressurealso tends to decrease F. However, when a consistent databaseis available (Zamora, 2000), it can be observed that high-Pmelts tend to be less calcic and somewhat more sodic thantheir lower-P counterparts (Moyen, 2011a and Fig. 10b). Thisis due to the lower Na2O and higher CaO contents in the high-pressure grossular+omphacite residual assemblage, comparedto the low-pressure amphibole+plagioclase one. Consequently,high pressure melts tend to be more trondhjemitic whilst lowpressure melts are more tonalitic.
4.3. Residual minerals
The composition of the residue of melting strongly controls thetrace element signature of the melts, and has therefore been investi-gated in some details. One of the most relevant findings is that garnetstability curves are fairly consistent, throughout the experiments andirrespective of H2O contents or source compositions. The garnet-incurve is positively sloped, from 9—10 kbar at 700 °C to 15 kbar at1100 °C (except in Zamora, 2000, where high Na2O pushes garnet-into pressures some 3–4 kbar higher at low temperature, probably asan effect of the increased stability of the albite component of plagio-clase). In contrast, amphibole and plagioclase stability are highly var-iable (Moyen and Stevens, 2006), although amphibole generallydisappears between 22 and 26 kbar, and is never found to be stable
40
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1200 700 800 900 1000 1100 1200
1200 700 800 900 1000 1100 1200
Arc basalt
P (
kbar
)
T (°C)
P (
kbar
)
T (°C)
esponds tomodelswith a tholeitic source,whereas on the right hand side a calc-alkaline sourcef the residual minerals duringmelting, and the trace elements composition of the source (and,esented as a function of its proportion in the total, partially molten system (i.e. melt+residue).
322 J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336
beyond 1100 °C. Plagioclase stability is primarily controlled by watercontents; in fluid-absent systems, it can be stable up to 1100 °Cwhereas fluid-present melting consumes it much more quickly (typ-ically before 950 °C at 10 kbar).
4.4. Trace elements: direct and indirect constrains
Central to the definition of TTGs is their trace element signature, asit has strong implications on the conditions of melting. However, di-rect determination of trace elements in experimental glasses hasonly recently become feasible, and previous studies had to rely onmore indirect approaches.
4.4.1. Indirect determination based on phase stabilityThe knowledge of both the modal composition of the residue and
the mineral/liquid partition coefficients, allows computing the traceelement concentration in experimental melts (Fig. 12). This approachis strongly dependent on the values selected for the partition coeffi-cients and, therefore, somewhat less precise than a direct analysis.However it allows modelling of the trace element composition of amelt over a large range of P–T conditions. (Moyen and Stevens,2006; Nagel et al., 2012; Nair and Chacko, 2008).
Such models have the potential to produce “geochemical” ther-mometers or barometers, allowing the interpretation of the composi-tion of igneous rocks in terms of a proxy for the PT conditions ofmelting. While such a quantitative approach is still beyond our capac-ities, an important qualitative conclusion is that HREE contents aswell as ratios such as Sr/Y or La/Yb are strongly pressure dependent;indeed, they are primarily linked to garnet abundance in the residue,itself being a function of pressure. Above the garnet-in line, the modalproportion of garnet progressively increases in the residual assem-blage from b5% at 10 kbar to up to 40% at 25 kbar, with matchingHREE depletions.
This does push the required depth of melting for “proper” (HREE-depleted) TTGs up to some 15 kbar at least (Bédard, 2006; Halla et al.,
Na 2
O
SiO2 S
Al 2
O3
Y
La/
Yb
60 65
23
45
67
8
60 65 70 75 80
1214
1618
0 1
050
100
150
200
250
0.05 0.10 0.20 0.50 1.00 2.00 5.00
0.5
1.0
2.0
5.0
1020
Ta
Nb
Nb/Ta=
5Nb/T
a=10
Nb/Ta=
20
Fig. 12. Different types of TTGs, identified by their geochemical features (Moyen, 2011a). Thpressure TTGs.
2009; Moyen, 2011a; Moyen and Stevens, 2006; Nair and Chacko,2008). It does also provide an explanation for the “high HREE” TTGs(technically not TTGs, based on our definition; Halla et al., 2009;Moyen, 2011a; Willbold et al., 2010), that could be generated at mod-erate depths of ca. 10–12 kbar, immediately above the garnet-in line,but still in the plagioclase stability field.
This line of evidence has been pushed further, with the identi-fication of several “sub-types” of TTGs (Halla et al., 2009; Moyen,2011a): the main differences reside in the HREE, Sr, and Nb–Tacontents (Fig. 10). Moyen (2011a) demonstrated that TTGs canbe classified in three groups: the “high pressure” group has lowHREE, low Nb and Ta, high Sr symptomatic of melts in equilibriumwith large amounts of garnet, some rutile, but no plagioclase. Incontrast, the “low pressure” group shows high HREE, Nb and Tabut lower Sr, consistent with plagioclase being stable in the resi-due of melting, but lesser amounts of garnet and no rutile. The“medium pressure” group is intermediate between the previoustwo. As plagioclase, rutile and garnet stabilities are stronglypressure-dependent; the three groups can be interpreted interms of depth of melting.
Therefore, both the experimental and the geochemical evi-dence indicate that the wide group of sodic Archaean granitoids(TTG s.l.) formed by melting over a large range of pressures. Nosingle model (i.e. eclogite vs. amphibolite residue, or shallow vs.deep melting) can account for the whole group, nor is applicableas a unique model for the generation of the Archaean continen-tal crust. Rather, it appears that diverse situations, probablycorresponding to contrasting tectonic scenarios, did exist in theArchaean.
4.4.2. Direct determination from in-situ analyses of experimental productsRecent analytical developments (LA-ICP-MS) have allowed the
analysis of small melt pools, and the direct determination of trace el-ement contents of experimental melts (Rapp et al., 2010). While thisapproach does not allow “mapping” of the PT space, it has confirmed
Sr
iO2 SiO2
La/Yb
Sr/
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b
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00
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e dark blue, light blue and green dots correspond respectively to high, medium and low
1
2
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SA
MP
LE/ C
HO
ND
RIT
E
La Ce Nd Sm Tb Yb LuEu LaCe Nd Sm Yb LuEu Dy Er
Bulk rock
8 Kbar, 1000°C
FSS D15
Bulk rock
10Kbar, 850°C
10Kbar,950°C
22 Kbar, 1050°C
32 Kbar, 1100°C 10 Kbar, 875°C
10 Kbar, 1000°C
a b
Fig. 13. Chondrite normalised (Masuda et al., 1973) REE patterns for liquids generated by experimental basalt melting a) source FSS=tholeiite; b) source D15=low-K tholeiiticamphibolite. When residual garnet is absent (green symbols) magma HREE contents remain similar to source; when garnet is present (dark blue symbols), magma becomes strong-ly depleted in HREE, resulting in TTG-like patterns (FSS is from Rapp et al., 1991 and D15 from Wolf and Wyllie, 1994).
323J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336
(Fig. 13) that significant HREE depletions and REE fractionation in liq-uids resulting from metabasalt melting do require garnet to be stablein sufficient amounts in the residue.
5. The 1990s to 2000s: link with adakites
Adakites were recognised in the early 1990s as a significant type ofmodern arc lavas (Defant and Drummond, 1990; Drummond andDefant, 1990), that show sufficient similarities to Archaean TTGs tobe regarded as possible modern equivalents. Therefore, unravellingthe petrogenesis of adakites can bring information on the origin ofArchaean TTGs (Martin, 1999).
Fig. 14. Various types of “Adakites” (modified from Moyen, 2009). The different sym-bols correspond to different rocks described as “adakite” or “adakitic” in the originalpublications. The yellow field shows “ordinary” arc rocks. The dark purple data pointsare the “high silica adakites” (HSA) of Martin et al. (2005), i.e. the lavas that do actuallymatch the original definition of adakites; light purple points are “low silica adakites”(LSA).
5.1. Adakite definition and the adakite–TTG comparison
Adakites, as originally defined (Defant and Drummond, 1990) are
“Andesites, dacites and sodic rhyolites (dacites being most commonproducts), or their intrusive equivalents (tonalites and trondhjemites)(…) characterized by >56% SiO2, >15% Al2O3 (rarely lower), usuallyb3% MgO (rarely above 6% MgO), low Y and HREE relative to island-arc andesites, dacites and rhyolites (for example, Y and Yb b18 and1.9 ppm respectively), high Sr relative to island-arc ADR (=andesites,dacites and rhyolites) (typically >400 ppm), low high-field strengthelements (HFSEs) as in most island arc ADR”.
This description can be compared to Martin's (1994) definition forthe TTG suite, presented in Section 2.2. Obviously, the two definitionsare very similar and the key points that distinguish TTG and adakitesfrom common granitoids or arc lavas respectively are the same, i.e.the intermediate to acid composition, with the lack of mafic compo-nents; the sodic nature; the high La and Sr, low Y and Yb contents,and correlated high Sr/Y and La/Yb ratios. Also implicit in the defini-tions of both adakite and TTG is the “arc” signature of either rockgroup, with a typical LILE/HFSE decoupling resulting in Nb–Ta nega-tive anomalies, low Nb/Th ratios, etc.; and a calc-alkaline nature. Itis worth pointing out that every aspect of the description above isan integral part of the definition of adakite (or TTG), and that norock should be called adakite (or TTG) if it does not match thewhole definition. There is, for instance, a disturbing trend to call“any rock with elevated Sr/Y or La/Yb ratios adakite” or “adakitic”, re-gardless of the other geochemical characteristics (see discussion inMoyen, 2009). Bearing this in mind, there is a clear, first order simi-larity between TTG and adakites (Martin, 1999; Martin et al., 2005),suggesting that both rock types formed by very similar processes.
The “conclusion” that adakites and TTGs formed under similarconditions has formed a strong basis for interpretations of the tecton-ic setting in which TTGs were generated. Since adakites (at least,modern adakites actually matching the definition outlined above)are found only in arc settings, they are a patent feature of subductionzone magmatism (irrespective of the details of the petrogenetic pro-cesses operating), and therefore it is tempting to conclude that TTGsare also subduction-related. Unfortunately, this logic has also been se-riously misused, mainly because a wide range of igneous rock hasbeen incorrectly described as “adakite” in the literature (Fig. 14;also see (Moyen, 2009)). This leads to spurious reasoning, wherebya rock, incorrectly classified as “adakite” is then demonstrated to beunrelated to subduction (Coldwell, 2008; Ding et al., 2011; Guo etal., 2007; Hastie et al., 2010; Li et al., 2011; Xu et al., 2006), leadingto the conclusion that “adakites are not related to subduction process-es” and therefore that “TTGs are not subduction-related either”. In the
Rb Ba Th U K Nb LaTa Ce Sr Nd HfZr Sm Ti DyGd Y Er Yb Sc V Cr Ni
0.1
0.01
1
10
100
RO
CK
/ P
RIM
ITIV
E M
AN
TL
ELSA
HSA
TTG
Fig. 15. Comparison of trace element contents in average TTG, High SiO2 Adakites (HSA) and Low SiO2 Adakites (LSA) showing that the HSA are very similar to Archaean TTG, but the LSA aresignificantly different. For instance, LSA display a positive Sr anomaly that does not exist in TTG andHSA, similarly LSA have higher concentrations of REE and transition element than TTG andHSA.
324 J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336
following discussion, we will, therefore, ignore all the “adakites” thatdo not correspond to the proper definition.
5.2. Are adakites and TTG really similar?
Following the initial discovery, in the 1990s, that adakites are con-vincingly similar to TTGs (Drummond and Defant, 1990; Martin,1999), this analogy was questioned and challenged. Both Smithies(2000) and Condie (2005b) pointed out the differences between thetwo rock types. In fact, based on geochemical considerations, Martinet al. (2005) discriminated two groups of adakites: “high-SiO2
adakites” (HSA) assumed to be generated by partial melting of a hy-drous basalt and “low-SiO2 adakites” (LSA) considered as partialmelts of a peridotite metasomatised by HSA. The two groups ofadakites differ by several characteristics, one of them being theirMg# (61 in LSA and 48 In HAS). The two groups also differ by their Niand Cr contents (in LSA Niaverage=103 ppm and Craverage=157 ppm;in HSA Niaverage=18 ppm and Craverage=26 ppm; Fig. 15). For TTGs,Mg# (average of 43), Ni and Cr contents (ca. 20 and 40 ppm respec-tively) are very similar to HSA, but strongly differ from LSA. Fig. 15also shows that Sr content is the same in both TTG and HSA(where this element does not correspond to any positive or negativeanomaly), while LSA display a strong positive anomaly, matchingmuch higher contents in LSA (Sraverage=2051 ppm) comparedwith both HSA and TTG (565 and 400–500 ppm respectively). Final-ly, the LREE content of LSA is greater than in both HSA and TTG. Con-clusively, it appears that most of the differences reported between
MgO %
5
7
6
4
3
2
1
040 50 60 70
T < 3 Ga
T > 3.5 Ga
T < 3.5 GaT > 3.0 Ga
Experimentalliquids
a
SiO2%
8
Fig. 16. TTG and experimental melts: MgO vs. SiO2 systematics. a) Comparison between experally more magnesian nature of TTGs compared to experiments. b) How mantle–melt intetiation trend (irrespective of the actual process generating it). Sample (1) has higher MgOsimply corresponds to a different position on the trend. Interactions of magma (2) with tha new trend (B). Sample (4) on this new trend actually has lower MgO than sample (2).(2)! This demonstrates that the relevant parameter is not the absolute MgO (or Ni, Cr, etc.)If the two yellow and green fields represent two sample sets, it is not immediately intuitiv
adakites and TTG come from the grouping of two distinct rock types,LSA and HSA, which are obviously different. There are, on the otherhand, no significant differences between HSA and TTGs (Martin et al.,2005).
6. The 2000s: TTG–mantle wedge interactions
Liquids produced by experimental melting of basalts appear tohave lower MgO (and Cr, Ni) than HSAs or TTGs at comparable SiO2
contents (Fig. 16). This feature is classically interpreted as reflectinginteractions between the primary adakite/TTG melt and the mantlewedge (Martin and Moyen, 2002; Martin et al., 2005; Maury et al.,1996; Prouteau et al., 2001; Rapp et al., 1999, 2010; Smithies, 2000;Stern and Killian, 1996). The TTG melt has a density lower thanthat of the surrounding rocks, such that it tends to ascend towardsthe surface. Consequently, if such a melt interacts with peridotite,this implies that its basaltic source is located under a peridotite(mantle) slice. Such a geometry is commonly achieved in subductionzone environments where oceanic lithosphere subducts under man-tle wedge (see Section 7 for further discussion on possible tectonicscenarios).
Conversely, the lack of clear geochemical evidence for melt–mantleinteractions (i.e., the relatively low Mg, Ni, Cr contents of some TTGs)has been used to infer that these rocks did not form under a mantleslice, thereby arguing against a “normal” subduction. Proposed alterna-tives include melting of the lower portion of a thick mafic crust(Atherton and Petford, 1993; Bédard et al., 2003; Condie, 2005b;
b
(A)(B)
SiO2%
(1)
(2)
(3)
(4)
MgO %
erimental melts and TTGs (grouped by age: Martin and Moyen, 2002) showing the gen-ractions would appear in Harker type diagrams, SiO2 vs. MgO. Trend (A) is a differen-than sample (2), but this does not reflect any sort of mantle interactions. Rather, it
e mantle will shift it to position (3); from this point onwards, the magma will defineYet samples (3) and (4) have experienced interactions with the mantle, not samplevalue, but the value at a given SiO2 content (i.e. the trend to which a sample belongs).e that the mantle-contaminated set is group (B).
325J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336
Hastie et al., 2010; Whalen et al., 2002); or a “flat subduction” scenario(Smithies, 2000; Smithies et al., 2003), with the subducted crust creep-ing under the upper plate without a proper mantle wedge (Gutscher etal., 2000).
It is, therefore, critical to understand the physical and petrologicalmechanisms of mantle–melt interactions, as well as their possiblegeochemical consequences in order to decide whether (some?) TTGsuites interacted with a peridotitic mantle — or not.
6.1. Mantle–melt interactions: evidence in space and time
Several lines of evidence have been used to infer mantle–melt in-teractions. Firstly, several groups (Condie, 2005b; Martin and Moyen,2002; Smithies, 2000) described a secular evolution pattern, withPalaeoarchaean TTGs having lower Mg#, Ni and Cr than their Neo-archaean counterparts. They interpret this evolution in terms of pro-gressively deeper basalt melting during the Archaean due to theprogressive cooling of the Earth. Indeed, a colder Earth would resultin lower geothermal gradients, such that slab melting would takeplace at greater depth and consequently under a thicker mantle peri-dotite slice, resulting in more efficient TTG melt–mantle peridotiteinteractions.
Since the middle of 1980s, plutons referred as “phenocryst-bearing”or as “Mg-rich granodiorites”were identified amongst the late Archaeanintrusions. Because the major element geochemistry of these rocks re-sembled that of Miocene high-Mg Andesite (Sanukite) from theSetouchi volcanic belt of Japan (e.g. Tatsumi and Ishizaka, 1982),Shirey and Hanson (1984) referred to them as “Archaean sanukitoids”.Thereafter, sanukitoids were described in almost every late Archaeancraton (Bédard, 1996; Champion and Smithies, 1999; Halla, 2005;Heilimo et al., 2010; Laurent et al., 2011; Leite et al., 2004; Lobach-Zhuchenko et al., 2005, 2008; Martin et al., 2010b; Moyen et al., 2001,2003; Samsonov et al., 2005; Smithies and Champion, 2000; Stern andHanson, 1991). Sanukitoids are monzodiorites to granodiorites, withan incompatible element signature similar to that of TTGs, but theyare more potassic and have higher compatible elements contents (fora given SiO2 content). Sanukitoids are similar to the “low silica adakites”(Martin et al., 2005). Based on these chemical features, a commonmodel for the origin of sanukitoids involves interactions between TTGmelts and the mantle (be it in the form of contamination of themagma during ascent, or of metasomatism and subsequent melting ofthe peridotite, Martin et al., 2010b), i.e. a process similar to the melt–mantle interactions discussed here,withmore pronounced and efficientinteractions.
Lastly, there is also physical evidence for primary adakite/TTGmelt andperidotite interactions. Adakitic (HSA) glassy inclusions have been de-scribed in olivine crystals from ultramafic mantle xenoliths in lavas fromthe Batan Islands, Philippines (Schiano et al., 1995). Interstitial adakiticglass is found in mantle xenoliths from the southern Andes (Kilian andStern, 2002), and MgO and Cr-enriched adakitic veins cut across mantlexenoliths in Kamchatka volcanoes (Kepezhinskas et al., 1995, 1996).
6.2. Petrological studies of TTG/mantle interactions and effects on themelt compositions
Most of the evidence presented above is indirect, and we haveonly limited understanding of the petrological or physical processesof melt–peridotite interactions, or of their effects on the melt'scomposition.
Firstly, the melt transfer mechanisms across the mantle are poorlyunderstood. Interactions would occur assuming that the melt perco-lates and permeates the mantle at the grain scale, therefore allowingfor important interactions. On the contrary, should melt ascend byfracturing and dyking, the scope of chemical interactions with thecountry rocks becomes very limited. Prouteau et al. (2001) also pro-posed that in this case the magma conduit will be armoured with
metasomatised peridotite, which will then become non-reactive.Studies of magma transfer in active margins suggests (i) a short trans-fer time from source to emplacement (Condomines et al., 2003;Hawkesworth et al., 2004; Jicha et al., 2007; Sigmarsson et al., 2002;Turner et al., 2003) and (ii) the focussing of magma transfer zonesin narrow regions with high magma fluxes (Grégoire et al., 2006;Rabinowicz et al., 1987, 2001). Both lines of evidence suggest thatthe opportunities for melt–mantle interactions may not be wide-spread, even in a subduction scenario.
Secondly, the petrologic (and geochemical) effects of such interac-tions on the melt are poorly known. While conventional logic sug-gests that the net effect might be an increase in compatible element(Mg, Cr, Ni) contents and related ratios (Mg#) in the melt, theamount of clear evidence for such an effect remains limited(Kepezhinskas et al., 1995, 1996; Rapp et al., 2010).
So far, only a few comprehensive studies were conducted on themelt–mantle reactions (e.g. Hoffer et al., 2008; Prouteau et al., 2001;Rapp et al., 1999, 2006, 2010; Sen and Dunn, 1994; Yaxley andGreen, 1998) and their effect on the melt's composition. Rapp et al.(1999, 2010) studied experimentally the origin of sanukitoids. Intheir experiments, the starting material consisted of TTG glass andcrushed peridotite. The TTG reacted with the peridotite by incongru-ent reactions forming orthopyroxene and amphibole, together with amelt enriched in compatible elements but relatively unaltered traceelements patterns and ratios. Geochemical modelling of such interac-tions (Moyen, 2009) leads to the same conclusion. Although thisstudy focussed on sanukitoids, it can be expected that less pro-nounced melt–mantle interactions (forming “contaminated” TTGs)would have similar, although more discreet, effects.
6.3. Compatible elements contents — an ambiguous marker?
Most of the geochemical evidence for mantle–melt interactionscomes from high Mg, Ni and Cr contents in TTGs, for which theremay be alternative explanations. For example, higher degrees of melt-ing of the same source will give more mafic melts. Likewise, mixingwith mafic magmas will have the same result. Entrainment of peri-tectic minerals (Stevens et al., 2007) will also yield a similar increasein compatible elements (compared to pristine melts). Lastly, highcontents should not be confused with elevated ratios (for instancehigh Mg#). A high Mg# can be achieved by increasing Mg, but alsoby decreasing Fe contents. Relatively leucocratic, high-Mg# meltscan be generated by high pressure melting (with Fe-rich omphaciticclinopyroxene as a residual phase (Laurie and Stevens, 2012; Moyen,2011a)); or by melt segregation and interactions with the continentalcrust during ascent (Getsinger et al., 2009). Mg# does therefore notcarry the same information than absolute MgO contents, and specifical-ly is no clear marker of the role of a mafic/ultramafic component.
Finally, it is worth stressing that making petrogenetic interpreta-tions based upon only a few elements is always hazardous. Indeed,MgO, Cr and Ni are very strongly negatively correlated to SiO2 inany magmatic series, simply reflecting chemical evolution duringdifferentiation. Therefore, comparing the MgO (for instance) con-tent of rocks at different SiO2 levels is uninformative and mislead-ing. This is illustrated in MgO vs. SiO2 Harker-type diagrams(Fig. 16b), where rocks belonging to a given magmatic series define“trends”. The position of a sample in a given trend is a function ofthe degree of differentiation (largely correlated to temperature,be it during melting or fractionation). Melt–mantle interactionswould be evidenced by samples “shifting” trend, i.e. moving to amore magnesian line of evolution. The absolute values on the otherhand would cover the same range in both cases. This demonstratesthat the relevant parameter is not the absolute MgO (or Ni, Cr, etc.)value, but the value at a given SiO2 content (i.e. the trend to which asample belongs).
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7. Geodynamic implications
An underlying question, present in all of the debate on TTG petro-genesis, is in what kind of geodynamic environment did these rocksform? From a pure geochemical point of view, TTGs are similar toarc rocks (although TTGs are notably depleted in HREEs and haveno negative Sr anomaly compared to typical arc lavas; Fig. 5b),which has been widely used as evidence for a subduction-related or-igin. However, the basic information provided by TTG geochemistry issimply that these magmatic rocks derive from melting of a hydrousmafic source. In addition, the depth of melting must be great enoughto stabilise garnet and sequester HREE — i.e., pressure must be >12–15 kbar (Moyen and Stevens, 2006). Consequently, any environmentallowing hydrous mafic rocks to melt at high pressure would beequally suitable to form “TTG-like” magmas, for instance an oceanicplateau (Willbold et al., 2009) or even a mid-oceanic ridge(Rollinson, 2009), if the mafic crust is thick enough. This purely com-parative approach is therefore a dead-end, and a deeper understand-ing of TTG genesis must be sought.
However, there is more than a simple qualitative link between thepressure of melting and the amount of garnet present in the residue(Moyen and Stevens, 2006; Nair and Chacko, 2008). Moyen (2011a)demonstrated that the TTG group is in fact made of several “sub-groups”, one of which does require large amounts of garnet, togetherwith rutile, in the residue. In fact the same geochemical effect can beachieved if garnet and rutile are phases fractionating during coolingand crystallisation (Macpherson et al., 2006), but the pressure re-quirements are the same for both models. In turn, this implies an evo-lution at pressures of about 20 kbar. This “high pressure TTG” group(or low HREE group in Halla et al., 2009) therefore requires the burialof hydrous mafic rocks into the mantle to depths in excess of 50–60 km, which is tantalisingly similar to subduction.
7.1. Plate tectonics in the Archaean?
The existence of plate tectonic processes in the Archaean is debat-ed, both on theoretical grounds and geological observation. The pri-mary cause for potentially different geodynamical behaviour in theArchaean is the change in thermal regime (Brown, 2006; van Hunenand van den Berg, 2008). The Archaean Earth produced two to threetimes the amount of today's radiogenic heat, and despite significantuncertainties in its early thermal history (e.g. Herzberg et al., 2010;Korenaga, 2006), the Archaean mantle was most likely hotter thanit is today, as evidenced by liquidus temperatures and MgO contentsof basaltic lavas through time (Abbott et al., 1994; Grove andParman, 2004; Herzberg et al., 2010; Jaupart et al., 2007), which sug-gests a secular cooling rate of about 100 K/Gyr. This translates into (i)more buoyant and more rapid and/or thicker oceanic crust (vanThienen et al., 2004), partly compensated by eclogitisation of buriedbasalt at depth (Korenaga, 2006; van Hunen and van den Berg,2008) as well as by the possible presence of komatiitic lavas associat-ed with oceanic basalts in the Archaean oceanic crust (Barbey andMartin, 1987); (ii) weaker lithosphere (as the material strength isstrongly temperature dependent); (iii) less efficient coupling be-tween the convective mantle and the overlying plates (O'Neill et al.,2007a, 2007b).
The net result on the convective style of Earth, and therefore onthe geodynamic style of the planet, is to weaken the efficiency of sub-duction. Possibly, the systemmay transition to a “stagnant lid” regime(Moresi and Solomatov, 1998), where the lithosphere (the conduc-tive boundary layer of the system) becomes separated from the con-vective mantle and “floats” on top of it without being recycled insubduction zones. Between plate tectonics and stagnant lid regimes,numerical models propose that the dominant tectonic style couldhave been intermittent (O'Neill et al., 2007a, 2007b), with the weakslab breaking off frequently (Barbey and Martin, 1987; Halla et al.,
2009; van Hunen and Moyen, 2012; van Hunen and van den Berg,2008) and therefore having little potential to pull the lithosphericplates. Thus the subduction dynamics would be different (DiGiuseppe et al., 2008; Sizova et al., 2010) with less coherent slabs(van Hunen and van den Berg, 2008), smaller scale and discontinuoussubduction. It is clear that if there was subduction during the Archae-an, its main features were significantly different, such that many ofthe geological markers used in recent terrains may well becomemeaningless.
Consequently, two main views are possible regarding Archaeantectonics. The first interpretation is that the Archaean tectonic styledid not involve any sort of plate tectonics— i.e., the thermal boundarylayer of the convecting system was not broken into plates integratedinto a mantle-scale convection system, through subduction zones andmid-oceanic spreading centres. Alternately, one may propose thatsome kind of plate tectonics operated — i.e., the thermal boundarylayer was mobile and recycled into the mantle along plate bound-aries, but that the shape, the size, the thermal conditions or/and thetiming of subductions were significantly different from the modernEarth. In addition, the Archaean aeon lasted 1.5 Ga and during this pe-riod, Earth cooled substantially, such that one could expect somechange in “plate tectonics” modalities through Archaean times.
Geological evidences for, or against, Archaean plate tectonics havebeen discussed in so many papers that a detailed account is beyondthe scope of this work. The existence of Archaean subduction zones(and, therefore, plate tectonics) is supported primarily by the exis-tence of a range of igneous rocks with an arc or arc-like compositionincluding boninites (Smithies et al., 2004), shoshonites (Kerrich andLudden, 2000), Nb-enriched basalts, and magnesian andesites andadakites (Polat and Kerrich, 2004). However it must be noted thatthe only requirement to form these rocks is the burial and dehydra-tion of surface rocks under a mantle wedge, irrespective of the actualsize and geometry of the buried portions of rock. Additional evidenceincludes the map pattern of many Archaean provinces (Czarnotaet al., 2010; Percival et al., 2002; Poujol et al., 2003), reminiscent oflateral accretion of stitched terranes (Coney et al., 1980), as well asthe existence of flat seismic reflectors, possibly representing fossilsubduction planes (de Wit and Tinker, 2004; Goleby et al., 2004;Ludden and Hynes, 2000). Finally, the existence of Archaean high-pressure and medium to low temperature metamorphism (Blocket al., 2012; Moyen et al., 2006; Saha et al., 2010, 2011; Volodichevet al., 2004), sometimes associated with “hotter” metamorphic rocks(Moyen et al., 2006; Stevens and Moyen, 2007) in Archaean ana-logues of paired metamorphic belts (Banno and Nakajima, 1992;Brown, 2002; Patrick and Day, 1995), suggests the existence ofsome form of subduction process in the Archaean.
In contrast, key indicators for modern plate tectonic processes (andin particular, subduction) are missing (Bédard et al., 2012; Hamilton,1998, 2003; Stern, 2005). No undisputed Archaean ophiolites areknown (Stern, 2007); andesites, blueschists and (in-situ) eclogites arealso missing (Stern, 2005). Well-characterised thrust-and-fold belts,accretionary wedges, and tectonic melanges are not known before ca.2.0 Ga. Rather, the dominant structures are best interpreted as bulkcoaxial shortening of the crust (Chardon et al., 2009) and gravity-driven tectonics involving the sinking of dense greenstone belts intothe soft basement (Bouhallier et al., 1995; Collins et al., 1998; Gormanet al., 1978), both suggesting deformation of a hot, “soft” Archaeanlithosphere (Choukroune et al., 1995) rather than plate boundaryprocesses.
7.2. TTGs as “non plate” magmas
Assuming the absence of plate tectonics in the Archaean, the pet-rological requirements to form TTGs can be accommodated in severalways. Indeed, the lack of plate tectonics does not imply a lack of man-tle convection; nor does it mean that mantle convection did not affect
Fig. 18. The “hot subduction”model (Martin, 1986; Martin et al., 2010b) for TTGs. a) P–T diagram showing the stability curve for hornblende (H), garnet (G) and rutile (R, greyband corresponding to the range of published estimates) as well as hydrous phases ofthe hydrated oceanic crust (A: anthophyllite; C: chlorite; Ta: Talc; Z: zoisite; Tr: trem-olite); the “wet” and “dry” solidus for basalts; and the typical along-slab geotherms inboth the modern and Archaean situation. The yellow field is the likely field for TTG gen-esis. b–d: cartoons corresponding to an early Archaean (b), late Archaean (c) and mod-ern (d) situation. During the Archaean, steepening geotherms result in progressivelydeeper melting of the slab; in the modern Earth, the slab cannot melt (panel a) but de-hydrates and yields common arc suites.. Same colour code and abbreviations as Fig. 17.
Fig. 17. Non-subduction models that fit the petrological constrains. a) Over a down-welling zone, portions of the lower crust (possibly thickened by tectonic stackingand/or plume activity) delaminate and sink into the mantle. Mafic rocks are thereforeheated and carried into the garnet stability field where they can melt. b) An upwellingmantle plume generates a thick oceanic plateau crust the base of which can melt togenerate TTGs. OC=oceanic crust, light green; CC=Continental crust, orange.Red=partially molten mafic rocks.Modified from Condie and Abbott (1999).
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the lithosphere. Rather, it means that the lithosphere was “stagnant”and did not break into pieces to be recycled into the mantle as part ofthe convecting system. However, mantle convection is able to affectthe lithospheric “lid” in at least two ways (Fig. 17).
Firstly, upwelling zones would cause mantle melting (in a scenariosimilar to modern plumes and oceanic plateaus), resulting in the ac-cumulation of large volumes of mafic magmas forming potentiallythick piles. Although the low viscosity of the lower crust will place alimit to the maximum thickness of these piles (England and Bickle,1984; Rey and Houseman, 2006), it is possible to accumulate some35–40 km of basalts, and therefore to reach the thickness requiredfor the formation of at least the low-pressure TTGs (TTG s.l., highHREE in Table 2). Melting of the base of the plateau crust (heatedby the mantle upwelling) would then generate TTG s.l. melts(Collins et al., 1998; Smithies et al., 2009; Van Kranendonk et al.,2007), that would be emplaced in a pile of dominantly “plume” relat-ed mafic to ultramafic lavas, similar to the situation observed in someArchaean cratons (and sometimes described as “plume–arc interac-tions”, Polat and Kerrich, 2001). This is neatly illustrated by the exam-ple of Kroksfjordur dacites in Iceland (Willbold et al., 2009). There,due to the interaction between the mid Atlantic ridge and a mantleplume, the crust can be 40 km thick (Kaban et al., 2002). However,despite their general similarity with TTGs s.l., the generated felsicmagmas are too shallowly formed to be in equilibrium with garnet,and therefore lack the TTG s.s. typical HREE depletion (Martin et al.,2010a; Willbold et al., 2010). Plateau melting thus cannot accountfor the genesis of “high pressure” TTGs (s.s.), requiring a much greaterdepth of melting. Finally, basalts at the base of a thick plateau-typepile are unlikely to become hydrated; therefore, the resulting meltswill, in general, not be TTG (or even I-type granites). This is illustratedby the fact that the typical felsic magma in any oceanic plateau is sy-enitic rather than I-type (Cousens et al., 2003; Gagnevin et al., 2003;Giret, 1990; Marsh et al., 1991; Martin et al., 2008; Shamberger andHammer, 2006). Consequently, melting of a thick basaltic plateaucan account for the origin of some TTG-like rocks (high HREE, or“low pressure” TTGs s.l.), but this model is not applicable to all TTGs(see Martin et al., 2008, for discussion).
Secondly, delamination at the mafic base of a crust could alsooccur, either above downwelling parts of the mantle convective sys-tem (Kröner and Layer, 1992); or as increasingly dense lower crustdetaches and sinks into the mantle. Increasing density can be a fea-ture of granulite-facies conditions; or, more efficiently, a result ofthe formation of an eclogite residue through a first stage melting ofthe lower crust (Bédard, 2006). In any case, fragments of maficrocks would sink into the mantle, where they can devolatilise (ifthey still have some water left), melt, and interact with the mantlein exactly the same way that subduction operates. Indeed, from a
petrological point of view, the processes operating are exactly thesame, whether in a true, large slab connected to an oceanic bottom;or in a small fragment of mafic rocks isolated in the mantle. Differ-ences could be found in terms of the size of the mafic body (andtherefore the rate of heating, the volume of magmas generated, etc.)but these aspects are not visible to petrology or geochemistry. Delam-ination is, therefore, able to generate magmas with the same “arc”signature that subduction produces.
7.3. TTGs as arc magmas
7.3.1. The hot subduction modelThe tectonic model that dominates the literature on TTG petro-
genesis, however, is probably the “hot subduction” model (Martin,1986; Fig. 18). Historically, this model is based on (i) the “arc” signa-ture of TTGs. However, this is a weak argument, as it is possible togenerate a “near TTG”, arc signature away from any arc (Rollinson,2009; Willbold et al., 2009). (ii) the requirement for melting a hy-drous basalt in the garnet stability field, i.e. >12 kbar. While moreconvincing than the previous argument this does not rule out the pos-sibility of melting at the base of the crust that could conceivably reacha thickness of 35–40 km. (iii) the need for pressures of ca. 20 kbar toform high pressure TTGs. Such pressures are extremely unlikely to beachieved within the crust, especially in an Archaean crust for whichthe higher heat flux from the mantle would result in a weaker crustunable to support a great thickness (England and Bickle, 1984; Reyand Houseman, 2006). Therefore, at least for the high pressure TTGs(that appear to represent maybe 20% of the Archaean TTG record,Moyen, 2011a), burial of mafic rocks from the surface seems to be arequirement; (iv) the necessity to have huge volumes of hydrous ba-salts at mantle depth; (v) the fact that present day HSA, essentiallysimilar to TTGs as discussed above, are found only in modern subduc-tion zones; and (vi) the existence of interactions between TTGmagmas (or “slab melts”) and mantle peridotite, likely in TTGs and
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Fig. 19. P–T diagram indicating the melting conditions of the three TTG groups ofMoyen (2011a) and the corresponding geotherms. The diagram shows linear geo-therms for reference, as well as the range of geotherms modelled for subductionzones, in purple; the solidus of amphibolites and multivariant amphibole (orclinopyroxene+epidote) breakdown reactions (red, shaded) together with the stabil-ity limits of garnet, plagioclase and rutile (grey field, range of published curves) inmafic compositions. Although melting occurs over a large range of temperatures (700to 1000 °C), this is still a relatively restricted range at the scale of this diagram, suchthat melting at different pressures as discussed previously implies melting on striking-ly different geotherms (discussion in text).
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obvious for sanukitoids (Section 6.1), demonstrates that melting oc-curred below a mantle slice.
7.3.2. Archaean vs. modern subduction systemsClearly, modern subduction processes do not generate large vol-
umes of TTGs or adakites. However, both geological and geodynamicconsiderations indicate that the differences between Archaean andmodern subduction processes are significant, such that perfect simi-larity should not be expected. In the subducting plate, higher mantletemperatures would favour slab melting over slab dehydration,thereby generating less andesites (related to slab dehydration), butmore TTG/adakites (slab melting) (Defant and Drummond, 1990;Martin, 1986). A possible analogue can be observed in certain arc seg-ments (such as the Austral Volcanic Zone in Southern Chile), wherethe subducting slab is young (or where the ridge itself is subducted),the magmas formed are adakites, rather than more typical calc-alkaline suites (Bourgois et al., 1996; Guivel et al., 1999; Martin,1999). In addition, common slab breakoff caused by thermal weaken-ing of the subducting lithosphere (van Hunen and Moyen, 2012; vanHunen and van den Berg, 2008) would mean that the buried litho-sphere was foundering and crumbling into the mantle rather thansubducting as long, continuous and persistent slab. This geometrywould result in short, discontinuous “bursts” of arc magmatism as op-posed to continuous arc activity, as now. Short-lived events of arc vol-canic and plutonic activity are indeed observed in the Archaean rockrecord (Moyen and van Hunen, 2012). In this view, the difference be-tween intermittent subduction and delamination would be minor,and largely transparent to petrology and geochemistry. A possiblysignificant petrological difference is the degree of hydration of the ba-saltic rocks in both scenarios; subduction of an oceanic floor will buryhydrated basalts, whereas lower crustal delamination will affect es-sentially dry rocks. Consequently, the temperature of melting, meltamount and residual assemblages produced (presence or absence ofamphibole) may also be expected to differ in both scenarios, andcould conceivably be tracked by investigating minute geochemicalor petrological differences.
7.4. Diverse geodynamic sites for the genesis of TTGs
There is a growing recognition of the fact that both subduction(slab melting) and intraplate models proposed above are probablyvalid petrogenetic processes for various rocks within the TTG s.l.group. Martin and Moyen (2002) proposed that the depth of hydrousbasalt melting changed in the course of Archaean times and Moyen(2011a) showed that the resulting TTGs (s.l; sodic granitoids) canbe subdivided into three sub-groups (see Section 3.3 and Fig. 12),corresponding to different depths of melting. Since the melting tem-perature is relatively constant regardless of pressure, this translatesinto melting along very different geothermal gradients (Fig. 19),from ca. 10–12 °C/km for the high-pressure group up to >30 °C/kmfor the low-pressure group. Regardless of the scenario preferred forArchaean tectonics, this is difficult to reconcile with a unique geo-dynamic setting, and it is therefore likely that TTGs s.l. can form in avariety of environments.
In Moyen's (2011) database, the high-pressure group (the mostlikely to be subduction-related) represents some 20% of the sodicgranitoids (corresponding to 10% of the grey gneisses); the low-pressure group (probably plateau-related; typically with HREE con-tents too high to match the TTG definition used in this work) corre-sponds to another 20% of sodic granitoids. The rest of the TTGs s.l.belong to the medium-pressure group and are somewhat ambiguous,as they seem to form along a 15–20 °C/km geotherm, too low for aplateau situation but probably too hot for subduction, even takinginto account the fact that Archaean subductions were likely hotter.In fact, they form along a geotherm similar to the one observed inthe Barberton region of south Africa at ca. 3.21 Ga (Nédélec et al., in
press), during the exhumation and melting of high-pressure meta-morphic rocks (Lana et al., 2010; Moyen et al., 2006). This suggeststhat the collapse of thickened orogenic (?) crust might be an optionto explore for the genesis of at least some portions of the TTG s.l.group. Indeed, Miocene adakite or adakitic rocks have been describedin Tibet, and related to melting of mafic lower crust in the thickenedcontinental lithosphere (Chung et al., 2003).
8. TTG through time
TTGs were originally identified as being a typically Archaeangroup of rocks. This contributed to the description of the Archaeanas a unique period in Earth history, with distinctive geological fea-tures. However, the growing interest in TTGs; the quest for modernanalogues of Archaean situations; the discovery of Hadean rocksand minerals; led to the progressive description of TTG (or TTG-like) rocks occurring during most of the geological record.
8.1. Early Archaean and Hadean TTG
The oldest rocks dated so far are the Acasta gneisses that outcropin the Northern Territories of Canada. They mainly consist of bandedtonalite and granodiorite whose zircons gave ages of 4.031±0.003 Ga(Bowring and Williams, 1999; Bowring et al., 1989). The major andtrace element composition of the Acasta gneisses is that of TTG s.s..They have low K2O/Na2O and fractionated REE patterns with low Ybcontent (YbN=3.6) (Bowring et al., 1990). The 4.0 Ga-old tonaliticdykes cutting across the Ujaraaluk unit (previously known as Fauxamphibolites, with a probable age of ca. 4.28 Ga: O'Neil et al., 2008)at Nuvvuagittuq, Canada are also TTG in composition (O'Neil, 2009).Recently discovered zircon cores from the Acasta gneisses weredated at 4.2 Ga (Iizuka et al., 2006; Iizuka et al., 2009). Their REE
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patterns are very similar to those of zircon extracted from Acasta TTG.Both lines of evidence demonstrate that 4.0–4.2 Ga ago the continen-tal crust already included TTG.
To date, the Hadean record is only accessible through zircon crys-tals. Most of these were discovered in Western Australia at Jack Hillsand Mont Narryer and are dated between 4.3 and 4.0 Ga (Cavosieet al., 2004; Compston and Pidgeon, 1986; Froude et al., 1983). Onlyone crystal, extracted from Jack Hills meta-quartzites gave an age of4.404±0.008 Ga (Wilde et al., 2001), which is the oldest age so farobtained on terrestrial material. The Jack Hills zircons contain severalmineral inclusions (quartz, plagioclase, K-feldspar, biotite, muscovite,etc., Cavosie et al., 2004; Hopkins et al., 2008, 2010; Maas et al., 1992;Menneken et al., 2007; Trail et al., 2004) characteristic of granitoids.Some of the Jack Hills zircons (Type 1 zircons of Hoskin, 2005) still re-cord a near-primary REE composition. The REE pattern of the magmafromwhich type-1 zircon crystallised (Fig. 20) can be calculated usingthe zircon/magma partition coefficients determined by Hinton andUpton (1991). These patterns are fractionated (high La/Yb) withstrong negative Eu and positive Ce anomalies, that Peck et al.(2001) ascribe to the unknown oxidation state during zirconcrystallisation. They probably do not reflect accurately the composi-tion of the parent magma. The first order feature remaining is, there-fore, the fractionated nature of the REE patterns of the liquids inwhich the Jack Hills zircons crystallised. Wilde et al. (2001) considerthat such REE patterns, together with the presence of muscovite andK-feldspar in the host magma, indicate that this latter must havebeen generated by recycling of older granitoids; the source of whichmust already have TTG-like fractionated patterns. Finally, an exhaus-tive isotopic Lu–Hf study (Blichert-Toft and Albarède, 2008) showedthat the magma from which the Jack Hills zircons crystallised had176Lu/177Hfb0.01. This is inconsistent with MORB, oceanic plateau ba-salts or arc magmas, but similar to Archaean TTG. The available evi-dence therefore suggests that Jack Hills zircons crystallised from aTTG-like granitoid melt (Blichert-Toft and Albarède, 2008; Guitreauet al., 2012; Nebel-Jacobsen et al., 2010). Consequently, TTG-likemagmas appear to be amongst the most ancient types of felsicmagmas on Earth, having existed since the early Hadean.
8.2. Proterozoic and Phanerozoic TTG
The end of Archaean aeon is characterised by a decrease in theabundance of TTG, progressively replaced by potassic I-type granit-oids. This shift in the magmatic record has been attributed to thechange from the archaic to modern style of plate tectonics (Martin,
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Fig. 20. Did Jack Hills zircons crystallise in a TTG magma? (a) REE patterns of the 4.40 Ga oldminerals preserved their primary magmatic composition (Hoskin, 2005). (b) REE patterns ozircon/melt partition coefficients (Hinton and Upton, 1991). It appears that this magma ha
1993; Stern, 2008) or to an episodic mantle overturn regime and pos-sible shutdown of plate tectonics at the close of the Archaean (Condieet al., 2009; O'Neill et al., 2007b; Silver and Behn, 2008). The latterwould be supported by the very low volumes of granitic magmatism(in general) between 2.45 and 2.2 Ga (Condie et al., 2009).
After 2.5 Ga, TTGs become increasingly scarce and subordinated.Some are described from the 2.2–2.1 Ga Trans‐Amazonian–Birimianorogeny (Almeida et al., 2007; Baratoux et al., 2011; Conceição deAraújo Pinho et al., 2011; Delor et al., 2003; Enjolvy, 2008); as wellas during the 0.8–0.6 Ga Pan‐African orogeny (Isseini, 2011). Theyare generally associated with huge volumes of either calc–alkaline ju-venile magmas (De Souza et al., 2007) or products of intracrustalmelting. Whether they are exact matches for the Archaean TTG s.s.as defined here, or are related but distinct, is still not clear.
The generation of TTG became less and less commonwith time, butnever totally stopped. Today, TTGs are restricted to exceptional geo-dynamic environments, most often in association with adakitic volca-nics. This is the case for the ca. 4 Ma–old Cabo Raper pluton, emplacedin the Taitao peninsula (Chile), in response of the Chile ridge subduc-tion (Lagabrielle et al., 1994, 2000). This TTG s.l. pluton (Bourgoiset al., 1996; Guivel et al., 1999) is associated with the adakites of theAustral Volcanic zone of the Andes (AVZ) (Sigmarsson et al., 1998).
9. Discussion, pending issues and perspectives
Many of the questions outlined above are still not totally resolved.Here we outline a few avenues that we regard as promising to explorefor a better understanding of TTGs.
9.1. The source of TTG magmas
Although most workers propose, explicitly or not, MORB-likesources for Archaean mafic rocks, several lines of evidence suggestthat enriched basalts are required. For instance the models of Martin(1987), described in Section 3.2, uses a sourcewith a slightly fractionat-ed REE pattern, like amodern E-MORB. An enrichedmafic source is alsorequired to account for the composition of TTGs in the East Pilbara(Smithies et al., 2009) and at Barberton (Moyen et al., 2007). Indeed,most Archaean mafic magmatic rocks show a somewhat enrichedtrace elements signature (Condie, 1981, 2005a; Hollings and Kerrich,2006; Jahn et al., 1980; Martin, 2011; Moyen, 2011b; van Hunen andMoyen, 2012). The dominant source of Archaean TTGs was probablynot MORBs, but rather an enriched tholeiitic basalt (Martin, 2011;Moyen, 2011b). It would be interesting to compare with more recent
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/ C
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ES
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zircon crystal (sample W74/2-36) from Jack Hills (Peck et al., 2001) showing that thesef the host magma of Jack Hills zircons, computed from the zircon composition and thed fractionated REE patterns like TTG.
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TTGs, and see whether this evolution mirrors an evolution in the com-position of the dominant type of mafic rocks.
l-P and I-type granitesl-P and I-type granitesl-P and I-type granites
Clinopyroxene entrainment trend
Clinopyroxene entrainment trend
Clinopyroxene entrainment trend
Garnet entrainment trend
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Garnet entrainment trend
Ca. 3.47 Ga Callina Supersuite (Pilbara)3.45 Ga Stolzburg Pluton (Barberton)3.21 Ga Kaap Valley Pluton (Barberton)
0.04 0.08 0.14
Fig. 21. Peraluminosity vs. “maficity” relationships (after Clemens et al., 2011). A/CNK:molar Al/(2Ca+Na+K); FM: molar Fe+Mg. In this diagram the leucogranites (felsicperaluminous, f-P of Villaseca et al., 1998) cluster on the low FM side of the diagramand define a weak vertical or no trend. The S-type granites (highly peraluminous, h-Pof Villaseca et al., 1998) plot in a positively correlated array. The I-types and “l-P”(low peraluminous, Villaseca et al., 1998) define a negative correlation (three colouredfields). The dashed arrows correspond to the trends of entrainment of peritectic min-erals (garnet or clinopyroxene) as proposed by Clemens et al. (2011). Three selectedTTG suites are plotted. The Stolzburg pluton defines a low-maficity cluster with aloose vertical correlation, whereas the Kaap Valley pluton and Callina supersuitesplot along negatively correlated arrays.
9.2. TTG with a long crustal residence?
Whole rock radiogenic isotopic studies in TTGs initially turned outto be rather disappointing. Despite some heterogeneities, the typicalTTG is near chondritic (εNd(T) ranging from −4 to +2 for instance),and no correlation can be demonstrated with other major or trace el-ement patterns. The slightly enriched elemental signature observedin at least some TTGs does suggest an enriched source Section 9.1),but this question has not been studied.
Recently, the advance of in-situ Hf (and O) isotopic studies in zir-cons has revived some interest for isotopic studies of TTGs (or greygneisses, described as TTGs in most of the works cited here and onemust bear in mind that several distinct rock types were thus probablyincluded under this name). A recent paper by Guitreau et al. (2012)reports >12,000 Lu–Hf analyses on zircons extracted from greygneisses, together with 141 Lu–Hf whole rock analyses on TTG s.l.samples. The age distribution in single zircons confirms the findingsof e.g. Albarède (1998), Condie (1998) and Rino et al. (2004). Theclustering of zircon ages suggests that crustal growth was episodicand occurred in three or four main episodes during the Archaean.Each episode of crustal growth starts with magmas having a chondrit-ic or positive εHf progressively, the zircon's isotopic signature evolvestowards negative εHf values. This evolution is interpreted in term ofrecycling of juvenile crust formed at the start of each “super-event”.
The 176Lu/177Hf ratio of juvenile granitoids formed during eachsuper-event did not vary throughout the Archaean, and remains closeto the chondritic value of 0.0336 (Bouvier et al., 2008). Consequently,it appears that since ~3.8 Ga the extraction of continental crust hasnot depleted its mantle source. A possible interpretation is that TTGsare generated from a mafic precursor ultimately extracted from theundepleted (lower)mantle (a possibility also supported by the trace el-ements data discussed in Section 9.1), such as basalts from oceanic pla-teaus formed above a deep-seatedmantle plume. The remaining plumeresidue would give rise to the depleted upper mantle.
The source of granitoids formed between the main events of crust-al growth has a long crustal residence time, possibly separated fromthe mantle several hundreds of Myr before the formation of the spe-cific rock specimen (Guitreau et al., 2010; Kröner et al., 2011; Zehet al., 2008, 2009). Furthermore, zircons extracted from gneissesfrom a specific region or terrane tend to follow a distinctive εHf vs.time evolution (Zeh et al., 2009), suggesting that all gneisses fromone area, regardless of age, ultimately derive from a similar mafic pre-cursor, extracted roughly at the same time from the mantle. Severalinterpretations can account for this observation: (i) the observed evo-lution could be dominantly a feature of recycling of older gneisses; inother words, the grey gneisses samples are actually dominated byrocks that are not TTGs but rather potassic granitoids formed by suc-cessive melting events of the earliest TTG gneisses. TTG would thenform only during the discreet, juvenile super-events. This corre-sponds to the evolution described in the East Pilbara (Championand Smithies, 2007), where successive granitic generations betweenca. 3.5 and 2.9 Ga become more potassic, but keep a similar Ndmodel age, suggesting that all the granitoids emplaced during ca.600 Myr formed by successive melting and refining of the earliestcontinental nucleus, made of juvenile TTGs intruded between 3.5and 3.4 Ga. (ii) Assuming that the sampled grey gneisses were actual-ly proper TTGs and were generated by melting of mafic rocks, thiswould imply that the mafic precursor was extracted from the mantlelong before it melted to yield TTGs. Different tectonic scenarios canallow such a model; for instance, extraction of thick mafic plateaus,that survive for a long period before being subducted (Guitreauet al., 2012; Martin, 2011; Moyen, 2011b) and melted to form TTGs
or a “stagnant lid” geology, where a persistent basaltic shell occasion-ally melted to form TTGs (Fig. 17).
In any case, it is difficult to see how the problem could be solvedwithout coupling all approaches and trying to reconstruct the natureof the source and its crustal residence time as well as the conditions ofmelting of the studied gneiss samples. Once again, it must be stressedthat naming a sample “TTG” is not enough to replace a careful petro-genetic study.
9.3. Petrology
Nearly all the discussion presented here, and indeed the best part ofthe literature, is based on geochemical considerations. Yet, TTGs aregranitoids (or meta-granitoids), and at least in some examples itwould be possible to use petrological information to constrain theirevolution. In particular, a salient feature of many TTG intrusions is thepresence of magmatic epidote (Bédard, 2003; Moyen et al., 2007). Ingranitic magmas, magmatic epidote crystallises at high pressure (>8–10 kbar, Schmidt and Poli, 2004; Schmidt and Thompson, 1996; Zen,1985), such that its presence places important constrains on petroge-netic and tectonic models.
Recent advances in granite petrology emphasised the role of themelting reactions and their products, both melt and peritectic min-erals (Clemens et al., 2009, 2011; Stevens et al., 2007; Villaros et al.,2009b) in shaping the geochemistry of granitic intrusions. In highlyperaluminous granites (=S-type granites s.s., Villaseca et al., 1998)A/CNK is positively correlated with Mg+Fe. This was interpreted byStevens et al. (2007) as reflecting the fact that S-type granites are amixture between melt and entrained peritectic garnet. In contrast,moderately or low peraluminous granites (Villaseca et al., 1998) aswell as metaluminous “I-type” granites (Clemens et al., 2011) showa negative correlation between A/CNK and Fe+Mg. Clemens et al.(2011) interpreted this correlation as evidence for the entrainmentof peritectic clinopyroxene together with the melt. Finally,leucogranites (felsic peraluminous in Villaseca et al., 1998) show nocorrelation; their compositions are reasonably similar to these oftrue experimental melts (Stevens et al., 2007). TTGs show either a
331J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336
negative correlation between A/CNK and Fe+Mg (Fig. 21) or for theleucocratic units (typically low HREE “high pressure” TTGs), no corre-lation. This is in good agreement with the low HREE content of TTGscompared to S-types: although S-type melts form in equilibriumwith garnet, the resulting granites are not HREE-depleted becauseperitectic garnet is formed together with the melt and entrainedwith it. In contrast, in TTGs garnet is not entrained, perhaps becauseit existed as a porphyroblast in the pre-melting metamorphic rocks,and was therefore not available for entrainment. Regardless of the de-tailed petrogenetic processes however, it is clear that the TTG groupencompasses rocks having contrasting relationships between A/CNKand Fe+Mg, and therefore having distinct petrogenesis, that shouldbe elucidated.
9.4. Improper uses, dubious comparisons and groupings
9.4.1. Definitions, classifications and terminologyAlthough the termof “TTG” is inwideuse, there is actually no formal,
official definition for this rock type. As the previous discussion has dem-onstrated, however, a large part of the debates surrounding the origin ofTTGs is probably tied to definition issues, with different scientist usingthe same name for different rocks. In particular, the geology of Archaeancratonic areas is commonly described in terms of a “basement” ofgneisses, intruded by plutons or overlaid by supracrustal greenstone se-quences. More or less implicitly, many geologists would regard thegneissic basement as a whole, made up of “TTG” gneisses, and thereforethe TTG acronym is frequently used as a synonym for “grey gneisses”.However, grey gneisses do not consist of TTGs alone, but rather includea range of granitoid (sodic and potassic) as well as miscellaneous (am-phibolites, leucosomes, restites, etc.) components. In Moyen's (2011)database, a mere 50% of the samples of Archaean grey gneisses wereclassified as TTG. Fig. 22 demonstrates that a significant portion of thepublished literature on “TTGs” is actually based on rocks that shouldnever have been classified as such.
9.4.2. TTGs are not lavas (≠adakites!)The “adakitic”model relies on the comparison of adakites, which are
lavas, with TTGs, which are plutonic rocks. One may, however, disputethe wisdom of comparing lavas and plutonic rocks exclusively on geo-chemical grounds. Firstly, whenever coeval lavas and intrusive are ob-served (Bachmann and Bergantz, 2004) they typically do not have thesame composition; suggesting that the apparent similarity might notbe so significant. Secondly, lavas and intrusive rocks do not have the
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Fig. 22. Grey gneisses include more than TTGs (modified fromMoyen, 2011a). The darkblue, light blue and green dots are TTGs proper (as in Fig. 10). The grey symbols (seeMoyen, 2011a for further details) are the other (non TTG) components of grey gneisses.
same behaviour, in terms of petrology. Lavas are erupted, meaningthat they do not cross their solidus on the ascent path, whereas plutonsrepresent magmas trapped at depth. Therefore, it is likely that the twotypes of magma have some significant differences, especially in termsofwater saturation (Clemens, 1998; Clemens andDroop, 1998). Thirdly,lavas (but typically not “high-SiO2 adakites”) do commonly belong todifferentiation trends, spanning a rather large range of chemical compo-sitions from basalts to intermediate or acid compositions. The same isnot true for plutonic rocks (and for TTGs in particular) where therange of composition in any given pluton or intrusive suite does typical-ly not exceed a few weight percent SiO2; demonstrating that while dif-ferentiation processes play an important role in shaping volcanic suites,they are not so significant for plutonic rocks. Finally, lavas are demon-strably pure or nearly pure liquids, with a subordinate amount of phe-nocrysts. The same does not apply to plutonic rocks, however, inwhich inheritance and transfer of solid crystals is common, includingaccessory minerals (inherited zircon cores), and, probably, also majorminerals carrying a significant portion of the budget of elements suchas Fe andMg (Stevens et al., 2007; Villaros et al., 2009a). The geochem-ical signature of plutonic and volcanic rocks, even if derived from thesame source, may therefore not be shaped by exactly the same process-es. Again, proper petrogenetic studies are required to address this issue.
9.4.3. TTGs — tonalites, trondhjemites and granodioritesFinally, this work has outlined the fact that TTGs (s.l.) are, actually, a
very diverse group (and grey gneisses even more so). Is there a uniqueTTG series, or does this term actually correspond to several rock typessharing similarities and typically tectonically mixed and interleaved?The evidence presented here does indeed suggest that the term “TTG” isa misnomer. There is no “TTG” series— there are tonalites (low/mediumpressuremelting ofmafic rocks), trondhjemites (high pressuremelting ofmafic rocks) and granodiorites (melting of enriched crustal lithologies).
10. Conclusion
Since their initial description, in the 1970s, TTGs became a matter ofconsiderable scientific interest. The discussion on TTG origin is con-nected to many important topics in Earth Science, such as the differen-tiation of the silicate Earth, the tectonic style of the Early Earth and itssecular evolution towards plate tectonics as we know it today, the ori-gin of arcmagmas in general (role of slabmelting, mantle–melt interac-tions), themeaning of differentiation “trends” in geochemical diagrams,etc. We now have a conceptual framework in which most researchersthink about TTGs. However, despite the large volume of literature avail-able, many first order questions still remain unanswered. Large parts ofthe controversies are probably the result of incorrect definitions ofwhatTTGs are, either in geochemical terms (low vs. high Al2O3 series, HREEcontents, K/Na ratios, etc.), or in terms of plain field descriptions (greygneiss complexes comprise more than TTGs, such that any studyattempting to unravel TTG geochemistry based on a “grey gneiss” sam-ple set is intrinsically flawed). After 40 years of advances in analyticaltechniques, with an increasing volume and diversity of data availableto the researcher, it isworth noting that old fashionedfield observationsand petrography still hold the key to the solution of some first orderproblems.
Acknowledgements
The authors are grateful to Hugh Smithies and Bor-Ming Jahn fordetailed and fruitful reviews which greatly improved the quality ofthis manuscript. Editorial assistance by Nelson Eby is also gratefullyacknowledged. This work evolved from a keynote talk given at the23rd Congress of African Geology, held in Johannesburg in January2011; the authors wish to thank the organizers of the conferencefor invitation. This review is the result of a long scientific associationwith TTGs and people trying to understand them, and we wish to
332 J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336
express our gratitude to all the colleagues who contributed material,data, ideas and questions over the years.
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