Forty years of TTG research

Lithos 148 (2012) 312–336

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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, France Clermont Université, Université Blaise Pascal, Laboratoire Magmas et Volcans, 5 rue Kessler, F-63038 Clermont-Ferrand Cedex, France c CNRS, UMR 6524, LMV, F-63038 Clermont-Ferrand, France d IRD, R 163, LMV, F-63038 Clermont-Ferrand, France b

a r t i c l e

i n f o

Article history: Received 31 October 2011 Accepted 8 June 2012 Available online 16 June 2012 Keywords: Archaean TTG Petrology Geochemistry

a b s t r a c t TTGs (tonalite–trondhjemite–granodiorite) are one of the archetypical lithologies of Archaean cratons. Since their original description in the 1970s, they have been the subject of many studies and discussions relating to Archaean geology. In this paper, we review the ideas, concepts and arguments brought forward in these 40 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. During this period, most work was focused on the identification and description of this suite, and the recognition that it is a typical Archaean lithology. TTGs were already recognised as generated by melting of mafic rocks. This was corroborated during the next decade, when detailed geochemical TTG studies allowed us to constrain their petrogenesis (melting of garnet-bearing metamafic rocks), and to conclude that they must have 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 experimental petrology work. A wealth of independent studies demonstrated that melting of amphibolites as well as of mafic eclogites can give rise to TTG liquids; whether amphibolitic or eclogitic conditions are more likely is still an ongoing debate. From 1990s onwards, one of the key questions became the comparison with modern 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 Archaean plutonic rocks, whether sodic or potassic, irrespective of their HREE contents. This leads to a large range of petrogenetic and tectonic scenarios; a fair number of which may well have operated concurrently, but are applicable only to some of the rocks lumped together in the ever-broadening TTG “bin”. © 2012 Elsevier B.V. All rights reserved.

Contents 1. 2.

3.

4.

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1970s: characterisation of TTG suites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.1. Field data and petrographic characteristics . . . . . . . . . . . . . . . . . . . . . . . 2.2. Geochemical characteristics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.3. Definition issues . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1980s: petrogenetic modelling based on geochemistry . . . . . . . . . . . . . . . . . . . . . 3.1. “Historical” models for TTG genesis . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.2. The reference model: melting of meta-mafic rocks . . . . . . . . . . . . . . . . . . . . 3.3. Melting of meta-basalts: at what depth? . . . . . . . . . . . . . . . . . . . . . . . . The 1990s: experimental petrology confirms that melting of mafic rocks in the garnet stability field 4.1. A review of experiments, starting materials and water availability . . . . . . . . . . . . 4.2. Major elements composition of melts . . . . . . . . . . . . . . . . . . . . . . . . . . 4.3. Residual minerals . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . gives rise to . . . . . . . . . . . . . . . . . .

⁎ Corresponding author at: Université Jean-Monnet, Université de Lyon, 23 rue du Docteur Michelon, 42023 Saint-Etienne, France. E-mail address: [email protected] (J.-F. Moyen). 0024-4937/$ – see front matter © 2012 Elsevier B.V. All rights reserved. doi:10.1016/j.lithos.2012.06.010

. . . . . . . . . . . . . . . . . . TTG . . . . . .

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4.4.

Trace elements: direct and indirect constrains . . . . . . . . . . . . . . . . . . 4.4.1. Indirect determination based on phase stability . . . . . . . . . . . . . 4.4.2. Direct determination from in-situ analyses of experimental products . . . . 5. The 1990s to 2000s: link with adakites . . . . . . . . . . . . . . . . . . . . . . . . . 5.1. Adakite definition and the adakite–TTG comparison . . . . . . . . . . . . . . . . 5.2. Are adakites and TTG really similar? . . . . . . . . . . . . . . . . . . . . . . . 6. The 2000s: TTG–mantle wedge interactions . . . . . . . . . . . . . . . . . . . . . . . 6.1. Mantle–melt interactions: evidence in space and time . . . . . . . . . . . . . . . 6.2. Petrological studies of TTG/mantle interactions and effects on the melt compositions 6.3. Compatible elements contents — an ambiguous marker? . . . . . . . . . . . . . 7. Geodynamic implications . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.1. Plate tectonics in the Archaean? . . . . . . . . . . . . . . . . . . . . . . . . . 7.2. TTGs as “non plate” magmas . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.3. TTGs as arc magmas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.3.1. The hot subduction model . . . . . . . . . . . . . . . . . . . . . . . 7.3.2. Archaean vs. modern subduction systems . . . . . . . . . . . . . . . . 7.4. Diverse geodynamic sites for the genesis of TTGs . . . . . . . . . . . . . . . . . 8. TTG through time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.1. Early Archaean and Hadean TTG . . . . . . . . . . . . . . . . . . . . . . . . . 8.2. Proterozoic and Phanerozoic TTG . . . . . . . . . . . . . . . . . . . . . . . . 9. Discussion, pending issues and perspectives . . . . . . . . . . . . . . . . . . . . . . . 9.1. The source of TTG magmas . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.2. TTG with a long crustal residence? . . . . . . . . . . . . . . . . . . . . . . . . 9.3. Petrology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.4. Improper uses, dubious comparisons and groupings . . . . . . . . . . . . . . . . 9.4.1. Definitions, classifications and terminology . . . . . . . . . . . . . . . . 9.4.2. TTGs are not lavas (≠adakites!) . . . . . . . . . . . . . . . . . . . . 9.4.3. TTGs — tonalites, trondhjemites and granodiorites . . . . . . . . . . . . 10. Conclusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

313

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42°

1. Introduction Barent

s Sea

68° KO

LA

Murmansk PE

NI

NS

UL

A

68° Kand alaks ha

Arctic Circle

W

hi

te

S

ea

Kemi

a

64°

lf

64°

of

Bo

th

ni

Onega

Gu

The first Archaean units investigated were greenstone belts, as they include contrasting lithologies, such as komatiites or banded iron formations, which are potentially ore-bearing and of economic interest. The more homogeneous and economically less promising granitoids remained almost unknown for a long period. Glikson (1979) stated that “The largely unknown nature of the granitic terrains is manifested by their presentation as undivided pink areas on regional geological maps and their designation in such terms as ‘sea of granite’ (Yilgarn, Western Australia) or ‘Peninsular Gneiss’ (Southern India)” (Fig. 1). Only at the end of the 1960s and beginning of the 1970s did geologists begin to show interest in the “sea of grey gneiss” (Anhausser et al., 1969; Bliss and Stidolph, 1969; Glikson and Sheraton, 1972; Heimlich, 1969; Heimlich and Banks, 1968; Hunter, 1970; Lund, 1956; McGregor, 1973; Sheraton, 1970; Viljoen and Viljoen, 1969). The development of geochronology, initially by K–Ar and Rb–Sr methods (Black et al., 1971; Goldich et al., 1970; Hanson et al., 1971; Heimlich and Banks, 1968; Kouvo and Tilton, 1966; Moorbath, 1975; Moorbath et al., 1972) led to the first absolute ages of the different components of Archaean granitoid basement. The first studies, for technical reasons such as sample freshness, or quality of the outcrops, focussed on late granites: typically potassic granites cutting across the surrounding gneisses. However, this was a major milestone, fixing reliable temporal markers for all Precambrian terrains — that is for about 88% of Earth history. At the same time, the petrographic and geochemical characterisation of the grey gneiss suites had been undertaken (Arth and Hanson, 1975; Barker et al., 1979; Bridgwater and Collerson, 1976; Condie and Hunter, 1976; Hanson and Goldich, 1972; Hunter et al., 1978; O'Nions and Pankhurst, 1978; Tarney et al., 1979; Weaver and Tarney, 1980, among the pioneers), revealing the close association of three (sodic) plutonic types: tonalites, trondhjemites and granodiorites. This characterisation led to use of the acronym TTG, first used in a publication by Jahn et al. (1981).

Pe

tro

zav

ods

k Lak eO ne ga

Lake Ladoga

Structural trends Phanerozoic

60°

Proterozoic High-grade terrane

Helsinki

60°

nla of Fi Gulf

24°

nd

Leningrad

Archaean

«Sea» of granite and gneiss Greenstone-belt

30°

36°

Fig. 1. A “sea of granites”: Archaean geology in the late 1970s. On this geological map of the Baltic shield [after Salop and Scheinmann (1969) and Gaál et al.(1978) in Condie (1981)], Archaean terrains were subdivided into two lithologies: greenstone belts and granito-gneissic basement. The latter was considered as a “sea” of granites and gneisses, without further distinction between these felsic lithologies.

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Since the 1980s, TTGs have been widely studied, with two main focuses: (1) understanding the tectonic regime of the Early Earth and (2) constraining the processes of differentiation of the continental crust. The aim of this paper is to review the history of concepts and questions relating to TTGs, broadly in chronological order. Many of these questions remain incompletely answered, and each section will therefore contain several questions and discussion points. At the end, we will outline a few further questions recently identified, that remain to be discussed.

a

2. 1970s: characterisation of TTG suites

Q

b

The pioneer studies on TTGs, in the late 1960s and 1970s, focussed on the description of these rocks, and their comparison with other common granitoids.

T 2.1. Field data and petrographic characteristics TTGs are the main components of Archaean terrains. They mostly crop out as diversely deformed plutonic rocks, ranging from pristine or nearly pristine plutons (Fig. 2a); to orthogneisses (Fig. 2b); to components of complex, banded gneiss complexes, often migmatitic and crosscut 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 differentiated rocks of the suite contain hornblende. The main accessory phases are allanite, pistacite, apatite, zircon, titanite and titanomagnetite; the presence of magmatic epidote is very significant, as it implies that crystallisation 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 calcalkaline differentiation trend (Lameyre and Bowden, 1982).

a

c

a

b A

A

c

P

Fig. 3. Modal composition of TTGs. a) In thin section, typical TTG consist of an assemblage of quartz + oligoclase + biotite, this later mineral carrying the metamorphic foliation (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 differentiation trend of Lameyre and Bowden (1982), (T = tholeiitic suite, A = alkaline suite; the calc-alkaline suites are subdivided into: a = K-poor; b = intermediate; c = K-rich).

b

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 homogeneous 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 Gneiss Complex (Swaziland).

Table 1 Proposed 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 TTG unless 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 plutonic rocks (low pressure “TTGs” in Moyen, 2011a).

Grey gneisses Plutonic component Grey gneisses - non granitoid portion

Sodic Potassic

TTG s.s. low HREE

med. HREE

high HREE

Polyphased orthogneisses - including enclaves, leucosomes, etc., but also a plutonic component General characteristics Plutonic rocks, possibly orthogneissified Amphibolites, (meta)diorites, (meta)tonalites, (meta)granodiorites, (meta)trondhjemites, (meta)granites

rare tonalites

Mineralogy

Tonalites and trondhjemites, rare granodiorites

Granodiorites and granites,

Trondhjemites common

Major minerals - felsic

Quartz + Oligoclase + K-feldspar

Major minerals - mafic

Biotite +- Hornblende

Tonalites common

Quartz + Oligoclase (An20-30 plagioclase). K-feldspar is rare Biotite + Hornblende +- magmatic epidote Hornblende commonly

Magmatic epidote common

abundant

Allanite, pistacite, apatite, zircon, sphene, titanomagnetite

Accessory minerals

45 < SiO2 < 80 % SIO2 > 64 %, commonly 70 % or greater 70 < SiO2 < 75 %

Degree of differentiation

68 < SiO2 < 72 %

Fe2O3*+ MgO + MnO + TiO2 <

Fe2O3*+ MgO + MnO + TiO2 < 5%

4% Major elements geochemistry

Variable Sodium/Potassium

K2O ca. 3%

systematics

3 < Na2O < 5 %

0,5 < K2O < 2 %

Aluminium

0,3 < K2O/Na2O < 0,6

> 14 % at 72% SiO2

> 16 % at 70% SiO2

> 15% at 70% SiO2

> 14 % at 70% SiO2

Yb < 4 ppm

Yb < 1 ppm

Yb < 1,5 ppm

Yb < 2 ppm

REE patterns

(La/Yb)N > 15 La > 25 ppm

La > 10 ppm Negative Nb-Ta and Ti anomalies

HFSE No Sr anomaly, weak negative LILE (and related)

4 < Na2O < 6 %

Na2O > 5%

0,8 < K2O/Na2O < 1

Trace elements geochemistry

65 < SiO2 < 70%

J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336

(meta-) Tonalites, trondhjemites, granodiorites, granites Rock types

No Sr nor Eu anomaly

Eu anomaly 5 < Sr/Y < 40

50 < Sr/Y < 500

20 < Sr/Y < 200

10 < Sr/Y < 50

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J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336

Anorthite

a

Tonalite Granodiorite

Trondhjemite Granite

Albite

Orthoclase K

b

2.3. Definition issues Although early definitions of TTGs were relatively unambiguous, the name has been used much more loosely in the literature. Tables 1 and 2 summarise the nature and composition of the various rock types that have been called “TTG”; in Table 1, the grey boxes highlight the features matching the definition of Sections 2.1 and 2.2. TTGs often occur as a component 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, the former contains a range of tectonically transposed components

CA Tdh

Na

Masuda et al. (1973) divided by 1.2). They also lack significant Eu and Sr anomalies (Fig. 5a, Martin, 1986) but do show negative Nb–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 ratios are high ((Zr/Sm)average = 44.3). The lack of both Eu and Sr anomalies associated with the low HREE content are interpreted as reflecting the presence of garnet and amphibole as well as the lack of plagioclase, either as residual or fractionating phases; such that, most often, a high-pressure origin has become implicit in the term ‘Archaean TTG’ (Champion and Smithies, 2003). In this respect too, this trace element signature makes TTG different from modern continental crust, which has both higher HREE contents (Ybaverage = 3.26 ppm), and negative Eu, Nb–Ta, Sr and Ti anomalies (Fig. 5b).

Ca

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) differs from that of the modern continental crust (yellow field); b) K–Na–Ca triangle contrasting the evolution of TTGs and modern calc-alkaline magmas (brown CA trend). CA trend is characterised by K-enrichment during differentiation, whereas TTGs show no clear trend and remain Na-rich and K-poor.

La Yb N

Archaean TTG

a

150

ROCK / CHONDRITES

316

100

Post 2.5 Ga granitoids

100

10

2

1 1

La Ce Nd SmEuGdTbDy Er Yb Lu La Ce Nd SmEuGdTbDy Er Yb Lu

2.2. Geochemical characteristics 50

0 0

ROCK / PRIMITIVE MANTLE

TTGs are silica-rich (SiO2 > 64 wt.%, but commonly ~70 wt.% or greater) with high Na2O contents (3.0–7.0 wt.% Na2O) and correlated low K2O/Na2O (b0.5). They are poor in ferromagnesian elements (Fe2O3* + MgO + MnO+ TiO2 ≤ 5 wt.%, Fe2O3* = total Fe expressed as Fe2O3), with an average Mg# of 0.43 and average Ni and Cr contents of 18 and 40 ppm respectively (Table 1). In the normative classification diagram for granitoids (Fig. 4a) (O'Connor, 1965), the plutonic component of Archaean grey gneisses all plot in the trondhjemite, tonalite and granodiorite domains, thus accounting for the classically used TTG acronym (Jahn et al., 1981). This is in sharp contrast with modern granitoids, which are commonly richer in K2O (granodiorites to granites). Similarly, as shown in the cationic K–Na–Ca triangle, granitoids from the modern continental crust evolve through K-enrichment during differentiation, while Archaean TTGs do not show a clear differentiation trend, but plot close to the Na apex (Fig. 4b). Based on Al2O3 content, Barker and Arth (1976) subdivided trondhjemites into high- and lowAl2O3 groups; most Archaean TTG having Al2O3 > 15% at SiO2 = 70%, belong to the high‐Al2O3 group. They evolve from metaluminous to slightly peraluminous (A/CNK = 1; normative corundum b 1%), with A/CNK increasing towards silica-richer compositions. TTGs exhibit characteristic trace element signatures, with high contents in light Rare Earth Element (LREE) (Laaverage = 31.4 ppm) but very low 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

4

8

12

16

20

(YbN)

100 10 1 0.1 0.01

b RbBaTh U K Nb Ta La Ce Sr Nd Zr Hf SmEu Gd Ti Dy Y Er Yb V Cr Ni

Fig. 5. Trace element features of TTGs. a) (La/Yb)N vs. YbN plot for Archaean TTG (blue points) and modern continental crust (yellow points). A clear compositional change appears at the Archaean–Proterozoic boundary (~2.5 Ga). Archaean TTG (Inset 1) have low-HREE contents (0.3 b YbN b 8.5) associated with strongly fractionated patterns (high (La/Yb)N), while the post-2.5 Ga continental crust has high-HREE contents (4.5 b YbN b 20) 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 modern continental crust (brown).

J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336

317

Table 2 Average composition of (1) Archaean grey gneisses, unsorted; (2) plutonic, plagioclase-dominated component of the grey gneisses; (3) potassic and (4) sodic (=TTG s.l.) plutonic components, the sodic TTGs being themselves subdivided into three groups. Authors' database (Martin and Moyen, 2002; Moyen, 2011a). Mg# = cationic ratio Mg/Mg + Fe; A/ EuN CNK = cationic ratio Al/2Ca + Na + K; (La/Yb)N = chondrite-normalised La/Yb; Eu=Eu
¼ pffiffiffiffiffiffiffiffiffiffiffiffiffiffi , the N subscript indicates values normalised to chondrites (Masuda et al., 1973). SmN :GdN

Major elements are given in weight% and trace elements in ppm. Avg = average (bold); std = standard deviation. Grey gneisses All grey gneisses (N = 5851)

Plutonic component All plutonic plagioclase-rich component (N = 1973)

Potassic (N = 392)

Sodic (TTG s.l.)

TTG s.s.

SiO2 TiO2 Al2O3 FeOt MnO MgO CaO Na2O K2O P2O5 Rb Ba Nb Ta Sr Zr Y Hf Ni Cr V U Th La Ce Nd Sm Eu Gd Dy Er Yb Lu K2O/Na2O Mg# A/CNK Sr/Y (La/Yb)N Eu/Eu* Lu/Hf Nb/Ta

All sodic (TTG s.l.) (N = 1439)

Low HREE (N = 376)

Med. HREE (N = 788)

High HREE (N = 184)

Avg

Stdev

Avg

Stdev

Avg

Stdev

Avg

Stdev

Avg

Stdev

Avg

Stdev

Avg

Stdev

67.67 0.44 14.87 3.58 0.07 1.86 3.35 4.18 2.60 0.16 83.98 717.29 8.40 0.90 455.56 162.61 15.82 4.73 50.02 98.94 56.56 1.96 10.93 37.59 72.39 29.87 5.07 1.20 4.42 3.48 1.68 1.37 0.23 0.62 0.48 0.94 28.8 18.5 0.77 0.048 9.3

7.16 0.35 1.99 3.04 0.10 2.92 2.52 1.22 1.54 0.20 64.12 700.52 7.36 1.38 334.26 116.62 17.22 3.35 137.43 279.71 67.42 2.52 14.12 36.76 70.02 31.11 5.22 1.03 6.31 5.13 1.66 1.52 0.23

70.04 0.33 15.19 2.49 0.05 0.99 2.75 4.64 2.25 0.11 76.69 568.57 6.58 0.91 438.28 145.70 11.77 4.27 30.61 38.01 30.28 1.92 9.90 29.47 56.93 21.40 3.70 0.92 3.89 3.17 1.24 1.05 0.18 0.48 0.42 1.01 37.2 18.8 0.74 0.042 7.3

3.41 0.16 1.17 1.20 0.08 0.71 1.22 0.87 1.24 0.08 56.39 406.14 5.14 1.31 228.04 82.42 14.66 2.01 83.51 63.98 22.57 2.38 16.78 22.14 43.95 17.49 2.90 0.69 9.11 7.60 1.24 1.29 0.19

72.61 0.26 14.07 1.91 0.04 0.54 1.64 3.90 3.92 0.09 145.31 649.44 11.13 1.36 224.26 173.87 22.52 5.32 11.57 29.89 22.38 2.91 20.67 44.52 85.82 31.40 5.47 1.11 7.94 6.12 2.06 2.05 0.31 1.01 0.33 1.03 10.0 14.6 0.51 0.058 8.2

2.55 0.15 0.93 0.90 0.03 0.46 0.80 0.82 1.15 0.07 65.98 454.30 6.39 1.79 204.55 96.63 16.85 2.33 9.35 43.37 17.67 2.78 29.72 29.35 55.77 24.44 3.99 1.12 16.76 12.81 1.57 2.05 0.26

69.15 0.36 15.53 2.73 0.05 1.16 3.14 4.84 1.70 0.12 55.55 530.64 5.20 0.68 492.91 139.51 9.18 4.00 36.35 40.36 32.86 1.42 5.72 24.73 47.15 18.16 3.03 0.84 2.33 1.70 0.85 0.71 0.12 0.35 0.43 1.00 53.7 23.6 0.97 0.031 7.6

3.29 0.16 1.05 1.21 0.08 0.72 1.12 0.77 0.71 0.07 32.15 379.76 3.59 0.94 203.42 73.49 13.01 1.83 94.06 69.29 23.33 1.86 4.46 16.43 33.10 12.82 1.96 0.38 1.92 1.56 0.83 0.61 0.11

71.12 0.25 15.55 1.83 0.04 0.79 2.63 5.24 1.71 0.09 46.44 541.81 2.91 0.52 583.32 113.60 5.44 3.01 86.24 42.49 22.75 0.83 3.86 16.35 28.72 11.64 1.82 0.59 1.24 0.84 0.41 0.38 0.07 0.33 0.43 1.02 107.2 28.7 1.19 0.022 5.6

2.38 0.11 1.02 0.88 0.13 0.46 1.03 0.91 0.74 0.05 28.10 469.72 2.02 1.30 197.02 84.55 17.86 1.02 162.46 50.49 17.48 0.82 3.35 10.34 19.14 8.65 0.61 0.17 0.44 0.39 0.21 0.19 0.04

68.29 0.40 15.65 3.02 0.05 1.33 3.44 4.70 1.64 0.14 54.98 531.58 5.25 0.81 483.31 143.23 8.31 4.17 21.56 38.16 36.48 1.31 6.16 26.88 52.87 19.31 3.18 0.91 2.43 1.63 0.81 0.70 0.13 0.35 0.44 0.99 58.1 26.0 1.01 0.030 6.5

3.20 0.15 1.02 1.09 0.04 0.73 1.08 0.65 0.61 0.08 29.54 348.96 2.41 0.98 172.07 62.77 5.21 1.84 52.98 81.95 24.36 1.58 4.24 14.06 29.06 9.29 1.20 0.30 0.90 0.77 0.41 0.34 0.09

68.81 0.42 15.21 3.31 0.06 1.29 3.24 4.47 1.76 0.13 70.58 446.62 8.07 0.79 327.65 173.53 18.24 4.51 15.21 27.59 41.82 1.83 7.16 31.03 57.91 22.51 3.76 0.95 3.15 2.68 1.39 1.18 0.19 0.39 0.41 1.00 18.0 17.7 0.84 0.043 10.3

3.53 0.16 1.22 1.35 0.06 0.83 1.12 0.60 0.73 0.06 39.25 311.56 3.18 0.39 159.06 70.59 18.44 1.27 13.43 30.48 26.19 1.96 4.89 17.33 32.22 13.26 1.65 0.32 1.21 0.94 0.57 0.55 0.09

(Fig. 2c, d), not all of them being metagranitoids. Most outcrops are heterogeneous and not only contain TTG, but also K-rich granites, leucosomes, restites, and even amphibolites or metapelites — all encompassed under the global description of “grey gneisses”. In addition, the plutonic components of grey gneisses actually include a wide range of rocks (Moyen, 2011a) that comprise both sodic and potassic igneous rocks — of which only the sodic components are TTGs, by any reasonable definition. Similar potassic plutonic rocks are also emplaced as individual plutons, sometimes referred to as “enriched TTGs” or “transitional TTGs”. Finally, even the sodic plutonic (or metaplutonic) rocks include both low- and high-Al2O3 compositions, typically correlated with more or less fractionated REE patterns. Strictly speaking, only the high-Al2O3 rocks with fractionated 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, we will restrict the use of the term “TTG” to the high-Al2O3 sodic plutonic or metaplutonic igneous rocks with fractionated REE patterns and low HREE contents (i.e. low and medium HREE sodic in Table 2). “TTG s.l.” will be used in a broader sense to refer to 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 petrogenetic scenarios were proposed to account for the geochemical characteristics of Archaean TTG.

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3.1. “Historical” models for TTG genesis

Sr/Y 200

The four main models proposed to account for the origin of TTGs were:

High-P frac. cryst. of arc basalt

50

Arc basalt melting

150

1) Fractional crystallisation of a wet basaltic magma leaving a cumulate made up of hornblende ± plagioclase ± biotite (Arth, 1979; Arth et al., 1978; Barker, 1979; Maaløe, 1982; Smith et al., 1983). However, this model requires more than 75% fractional crystallisation to generate trondhjemitic liquids from a basaltic source and should therefore 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 trends from the parental (basalt) to the differentiated liquid, and such trends are not observed in TTG suites. Finally, a fractionation model requires concomitant basaltic and TTG magmatism, both in time and space. Although Archaean cratons do include significant amounts of mafic rocks (in the form of lavas from greenstone belts), detailed geochronology demonstrates that the TTGs and the basalts are seldom coeval. In fact, TTG emplacement is concomitant with breaks in the mafic activity in the nearby greenstone belts, during which the erupted lavas are rather felsic and similar to the TTG plutons (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 a compositionally similar group of rocks called adakites (see Section 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 and Scaillet, 2003) fractional crystallisation was proposed to play a role in shaping the geochemistry of some arc suites (Richards and Kerrich, 2007). While amphibole has only a limited potential to create the typical HREE depletion observed in adakites/TTGs (Davidson et al., 2007), garnet fractionation is a much more significant process that has been demonstrated both geochemically (Macpherson et al., 2006) and experimentally (Alonso-Perez et al., 2009) to be capable of generating HREE-depleted intermediate magmas from an ordinary andesitic parent at pressures of ca. 10 kbar (Fig. 6). While this process is undoubtedly able to generate some high Sr/Y or La/Yb magmas (Moyen, 2009), its applicability to TTGs is somewhat debatable. Firstly TTGs as well as adakites do not show long differentiation trends and the “less differentiated” (basalt to basaltic andesite) endmembers are typically missing. Secondly, an important feature of both adakites and TTGs is that their K/Na ratio remains low throughout differentiation, and does not correlate with differentiation indicators such as SiO2. On the contrary, whenever fractionation of amphibole or garnet is suspected (Macpherson et al., 2006; Richards and Kerrich, 2007) or experimentally observed (AlonsoPerez et al., 2009), the melts invariably become more potassic (also see Fig. 14 below) in the course of differentiation. Finally, in general one may dispute the similarity between plutonic and volcanic rocks, the latter being typically dryer magmas, with more potential to ascend, erupt, and undergo differentiation (Clemens and Droop, 1998). 2) Direct melting of the mantle, possibly metasomatised by fluids (Moorbath, 1975; Peterman and Barker, 1976). In this case very low degrees of melting (b5%) are required to generate felsic magmas. However, even at such low degrees of melting, theoretical calculations based on REE concentrations demonstrate that it is impossible to account for the Yb depletion and the high La/Yb ratios typical of Archaean TTG (Fig. 7 and Jahn et al., 1981, 1984; Martin, 1987; Martin et al., 1983). Furthermore, a quartz-normative magma, such as TTG, is unlikely to be in equilibrium with ultramafic mantle. 3) Partial melting of Archaean greywackes (Arth and Hanson, 1975). Although this model of recycling explains the REE and

50

Slab melting

a

100

50 70

50

Mc Pherson et al. field of adakites

80 80 Defant & Drummond’s field of adakites

0 0

10

20

30

Y

Sr/Y 500 Defant & Drummond’s field of adakites

400

Pleistocene Pliocene Paleogene

b

300 Limit of fig. 6a

200

100

0

0

10

20

30

40

50

Y

Fig. 6. Fractionation of common mafic parents to give high Sr/Y melts (Macpherson et al., 2006). a) Various models for the origin of Tertiary adakites in Mindanao (Philippines). Slab melting, 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 viable models. An isotopic mismatch between the slab and the adakites, as well as a correlation between SiO2 and Sr/Y, are taken as further evidence favouring fractionation over a slab melting model (MacPherson et al., 2006). Note, however, the difference between the adakitic 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 high Sr/Y values observed in many other modern adakites (grey field in panel b).

some trace-element characteristics of Archaean TTG, it cannot account for major- and other trace-element behaviour (Martin, 1994). The sodic nature of TTGs in particular is at odds with this model. 4) Partial melting of a hydrated basalt metamorphosed at high pressure 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; Hunter et 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 accepted in the late 80s, with the further requirement that garnet must be stable in the residue in order to account for the HREE depletion of TTG, putting strong constrains on the depth of melting. Fig. 7 shows 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 TTG genesis had progressively evolved and matured into a 3-stage mechanism (Fig. 8):

J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336

Calculated magma composition

ROCK / CHONDRITES

100

150 125 100

E G25

1

F = 1%

100

50

F = 25%

10 5

Source = Pm Primitive Mantle

25

TTG

Calculated magma composition

F = 10%

ROCK / CHONDRITES

La Yb N

319

2

F = 50%

Calculated magma composition

F = 10%

3

F = 50%

50 Source = G 0 Archaean tholeiite

Source = G 25 Archaean tholeiite

10 TTG

5

TTG

10 1 La Ce

Nd

SmEu Gd Tb Dy

Er

Yb Lu

1

G10

75

a

50 10

25

Mm Pm

25 25

Nd

SmEu Gd Tb Dy

Yb Lu

La Ce

Nd

SmEu Gd Tb Dy

Er

Yb Lu

b 10

20 30

25

10

L2 20

G0

25

50

10 25 50

50

L3

30

10

4

Er

25

L1

0 0

La Ce

50

8

12

16

(YbN)

4

8

AT 12

16

(YbN)

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 metasomatised (Mm) mantle compositions correspond to melt in equilibrium with a lherzolite residue containing 10% (L1) and 5% (L2) garnet or 5% spinel (L3). Dotted lines are olivine fractional crystallisation trends at depth. Numbers indicate the degree of melting or of fractional crystallisation. Blue field=Archaean TTG. Inset (1) represents the modelled REE patterns for melting of a chondritic mantle source (Martin, 1987). Grey pattern=Primitive mantle (Pm); blue pattern=average TTG; pale yellow field=domain of REE patterns for liquids generated by 1 to 25% melting of the source. (b) (La/Yb)N vs. YbN diagram summarising the different models 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 as well as into eclogite (E). Blue field=Archaean TTG. Insets represents the REE patterns modelled for G0 (inset (2)) and G25 (inset (3)) (Martin, 1987). Grey pattern=Archaean tholeiite; blue pattern=TTG; pale blue and red field=domain of REE patterns for liquids generated by 10 to 50% melting of the source.

1) partial melting of the mantle to generate a basalt; 2) melting of this basalt, metamorphosed to garnet bearing amphibolite or eclogite, to give rise to TTG parental magma; 3) limited low pressure (intacrustal) fractional crystallisation to produce the differentiated TTG suites. The third step – fractional crystallisation – may or may not have affected all suites. Where it has occurred, it was in the form of extraction of hornblende ± plagioclase from the parental TTG magma. In all cases, the degree of fractionation was lower than 25% (Martin, 1987; Moyen et al., 2007); however, it is important in shaping the overall geochemical trend of a given pluton or suite. 3.3. Melting of meta-basalts: at what depth?

high Zr/Sm ratios in Archaean TTG preclude an eclogitic residue but are consistent with a garnet bearing amphibolite, assuming the source of TTGs had a chondritic Nb/Ta. Based on the same data, but with a low Nb/Ta basaltic source (not uncommon in the Archaean record), Rapp et al. (2003) arrived at the opposite conclusion (Fig. 9). We shall return to this question, after reviewing the results of experimental petrology studies. 4. The 1990s: experimental petrology confirms that melting of mafic rocks in the garnet stability field gives rise to TTG melts Following the consensus on the origin of TTGs by partial melting of meta-basalts, a large number of experiments were conducted from the 1990s onwards to investigate the feasibility of this scenario. A relatively recent paper (Moyen and Stevens, 2006) recorded in excess of

Later refinement of the general model persisted for the next two decades. For example, while it is widely accepted that TTGs were generated by melting of hydrous meta-basalt, the depth of melting remained hotly debated, with either a garnet–amphibolite or an eclogite residue appearing to be realistic models. Foley et al. (2002) investigated the distribution of HFSE in TTG liquids and concluded that the low Nb/Ta and

40

Source (Foley et al., 2002)

Nb / Ta

MANTLE

Stage 1

Rutile-eclogite

30

20

Eclogite

17 17

PM

Stage 2

Amphibolite

10

THOLEIITE

PM Residue Hbl+Grt+Cpx+Ilm±Pl

Source AB-1 (Rapp et al., 2003)

0

TONALITIC MAGMA

1

Stage 3

FC Cumulate Hbl+Ilm±Pl

T.T.G. SUITE

Fig. 8. Schematic diagram summarising the succession of the different mechanisms implied in Archaean TTG genesis (after Martin, 1993). Hbl = hornblende; Grt = garnet; Pl =plagioclase; Cpx= clinopyroxene; Ilm=ilmenite; PM=partial melting; FC=fractional crystallisation.

5

10

25

50

100

500 1000

Zr / Sm Fig. 9. Nb/Ta vs. Zr/Sm diagram comparing TTG composition (blue circles) with modelled composition. The red fields are from Foley et al. (2002), for melting of rutile bearing 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) by eclogite melting.

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J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336

to ca. 60%; K2O from 0.1 to 1.8% and Na2O from 1 to 4.3%. Mg# spread between 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 on the genesis of TTGs/adakites were mostly conducted in the garnet stability field, i.e. greater than ca. 10 kbar. Nevertheless, the P–T space has not been completely explored, as most studies concentrated on a positively sloped “band” from 700–900 °C at P b 10 kbar to 1000– 1100 °C at P > 25 kbar. This implies that the behaviour of the system is not described in some “exotic” parts of P–T space, such as the eclogitic region. Only a handful of recent studies tried to address this question (e.g. Laurie and Stevens, 2012; Skjerlie and PatiñoDouce, 2002). Finally, different situations relative to water availability have been explored (Moyen and Stevens, 2006; Vielzeuf and Schmidt, 2001): (1) Fluid-present melting, with water either as a free phase (added to the starting material) or released during prograde heating by breakdown of hydrous phases such as chlorite or epidote. Depending on the total amount of water initially present, the melt may or may not have the potential to dissolve all of it, such that the system can become fluid-absent after the first melting increments. (2) Fluid-

300 experiments in some 15 studies (Beard and Lofgren, 1991; PatiñoDouce and Beard, 1995; Rapp and Watson, 1995; Rapp et al., 1991; Rushmer, 1991; Sen and Dunn, 1994; Sisson et al., 2005; Skjerlie and Patiño-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 basaltic composition, ranging from basalts to basaltic andesites and generally belonging to the tholeiitic series, or being close to the calcalkaline/tholeiitic boundary. Nevertheless, there are fairly significant dissimilarities between the different starting materials, in terms of modal composition, bulk rock chemistry and mineral chemistry. Amphibole/plagioclase ratios vary from 0.18 to >4; quartz abundances in experimental sources ranged from zero up to 24%. Other minerals (chlorite, epidote, Ti-oxides, titanite) were described in some starting materials. Accordingly, SiO2 content of the sources ranged from ca. 47

K

K

Source and water availability

Pressure

Quartz-present, water sat. Quartz-present, “dry”

> 20 kbar 10 - 20 kbar < 10 kbar

Quartz-absent, water sat. Quartz-absent, “dry”

a

b

Na

Ca

Na

Tholeiitic basalt

Arc basalt

40

40

c

35

30

30

25

25

P (kbar)

P (kbar)

35

Ca

20 15

20 15

10

10

5

5

0 700

800

900

1000

1100

1200

d

0

700

800

900

1000

1100

1200

T (°C)

T (°C) Tonalite Trondhejmite Granodiorite

Solidus Plag out Amp out Gt in

Granite Fig. 10. Main compositional features of experimental melts of amphibolites. a) In a K–Na–Ca triangle, most experimental melts (colour-coded based on the presence or absence of quartz in the source and the water availability, fluid-present or ‐absent) match the composition of TTGs (light blue field in the background). Increasing F values result in evolution towards more calcic and alkali-poorer compositions (arrows). b) Focussing on one single set of experiments (with the same source and H2O contents), a pressure effect can be seen with high pressure melts being more sodic, and low-pressure melts more calcic (Data from Zamora, 2000; plotted in Moyen, 2011a). c) and d) are “maps” of the melts composition in the P–T space (Moyen and Stevens, 2006) for two groups of sources, tholeiitic (c, left) and arc-like (d, right). The coloured fields are modelled and the dots correspond to actual experiments. Three main controls can be observed on these diagrams: the role of the source (enriched source give more granitic/granodioritic compositions), the role of temperature (high F, “hot” melts are more tonalitic) and the role of pressure (high pressure melts are more trondhjemitic).

J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336

absent melting, with all water in the system accommodated in hydrous minerals (generally amphibole), whose destabilisation triggers melting (“dehydration melting”). This is the most commonly investigated scenario, as the general consensus is that fluid-present melting at >10 kbar (i.e., garnet stability field) is unlikely in nature; see however Laurie and Stevens (2012) for an alternative view. (3) Totally dry systems, with no water at all even in hydrous minerals. This scenario might be relevant to melting of a dehydrated eclogitic source, or to a lower crustal granulite, and has hardly been investigated (Springer and Seck, 1997). 4.2. Major elements composition of melts Despite the large range of sources and conditions, the results of all experiments are fairly consistent. The position of the solidus was found to be controlled primarily by water availability (Moyen and Stevens, 2006), while the melt fractions are controlled by the nature of the source and the amount of water in the system, with wet and quartz-rich sources being more fertile. Since melting is essentially a eutectic process, the composition of all experimental melts (regardless of the source composition) is rather similar. Broadly speaking, all melts (at reasonable temperatures) are leucocratic and sodic, and therefore very comparable to TTGs. In detail, the following factors seem to affect melt compositions (Fig. 11): (1) The degree of melting (F) is the most important parameter. As F increases, the melts become more mafic and calcic and K2O contents decrease. The behaviour of Na2O and Al2O3 depends on the nature of the source, these two elements being compatible in plagioclase-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 F can result from high temperatures, but also from water or quartz-rich sources. Therefore melting of dry meta-basalts tends

321

to yield granitic, rather than trondhjemitic melts, as the lack of water inhibits melting (Prouteau, 1999). (2) The composition of the source, and especially its K2O contents, is also significant. Indeed, in a system where no potassic mineral is present, K2O behaves as an incompatible trace element, such that its content in the melt is a function of both the source composition and the degree of melting (Shaw, 1970). More potassic sources thus yield potassic melts, granites and granodiorites rather than trondhjemites and tonalites (Martin and Sigmarsson, 2007; Sisson et al., 2005). (3) Pressure exerts a marginal control. This effect is difficult to identify, as it is hidden by the fact that increasing pressure also tends to decrease F. However, when a consistent database is available (Zamora, 2000), it can be observed that high-P melts tend to be less calcic and somewhat more sodic than their lower-P counterparts (Moyen, 2011a and Fig. 10b). This is due to the lower Na2O and higher CaO contents in the highpressure grossular+ omphacite residual assemblage, compared to the low-pressure amphibole + plagioclase one. Consequently, high pressure melts tend to be more trondhjemitic whilst low pressure melts are more tonalitic. 4.3. Residual minerals The composition of the residue of melting strongly controls the trace element signature of the melts, and has therefore been investigated in some details. One of the most relevant findings is that garnet stability curves are fairly consistent, throughout the experiments and irrespective of H2O contents or source compositions. The garnet-in curve is positively sloped, from 9—10 kbar at 700 °C to 15 kbar at 1100 °C (except in Zamora, 2000, where high Na2O pushes garnet-in to pressures some 3–4 kbar higher at low temperature, probably as an effect of the increased stability of the albite component of plagioclase). In contrast, amphibole and plagioclase stability are highly variable (Moyen and Stevens, 2006), although amphibole generally disappears between 22 and 26 kbar, and is never found to be stable

Solidus Plag out Amp out Gt in

Tholeiitic basalt

0.20 0.40

30

30 20 10 0

Arc basalt

40

P (kbar)

0.00

P (kbar)

Garnet

40

20 10 0

700

800

900

1000 1100 1200

700

800

T (°C)

100 150

40

40

30

30

P (kbar)

50

P (kbar)

(La/Yb)N

0

20 10 0

900

1000 1100 1200

T (°C)

20 10

700

800

900

1000 1100 1200

T (°C)

0

700

800

900

1000 1100 1200

T (°C)

Fig. 11. Mapping of garnet abundance (top) and La/Yb (bottom) in the PT space. The left column corresponds to models with a tholeitic source, whereas on the right hand side a calc-alkaline source with an arc-like trace elements pattern is used; this affects both the stability fields and abundance of the residual minerals during melting, and the trace elements composition of the source (and, therefore, of the melts). See Moyen and Stevens (2006) for further details. Garnet abundance is represented as a function of its proportion in the total, partially molten system (i.e. melt+residue).

J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336

beyond 1100 °C. Plagioclase stability is primarily controlled by water contents; in fluid-absent systems, it can be stable up to 1100 °C whereas fluid-present melting consumes it much more quickly (typically 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, as it has strong implications on the conditions of melting. However, direct determination of trace elements in experimental glasses has only recently become feasible, and previous studies had to rely on more indirect approaches.

4.4.2. Direct determination from in-situ analyses of experimental products Recent analytical developments (LA-ICP-MS) have allowed the analysis of small melt pools, and the direct determination of trace element contents of experimental melts (Rapp et al., 2010). While this approach does not allow “mapping” of the PT space, it has confirmed

Sr

5

Na2O

3

4

14

65

70

75

80

60

65

a=

75

80

Sr/Y

1

50 0

0.05 0.10 0.20 0.50 1.00 2.00 5.00

70

5 10

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4.4.1. Indirect determination based on phase stability The knowledge of both the modal composition of the residue and the mineral/liquid partition coefficients, allows computing the trace element concentration in experimental melts (Fig. 12). This approach is strongly dependent on the values selected for the partition coefficients and, therefore, somewhat less precise than a direct analysis. However it allows modelling of the trace element composition of a melt 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” thermometers or barometers, allowing the interpretation of the composition of igneous rocks in terms of a proxy for the PT conditions of melting. While such a quantitative approach is still beyond our capacities, an important qualitative conclusion is that HREE contents as well 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 modal proportion of garnet progressively increases in the residual assemblage from b5% at 10 kbar to up to 40% at 25 kbar, with matching HREE depletions. This does push the required depth of melting for “proper” (HREEdepleted) TTGs up to some 15 kbar at least (Bédard, 2006; Halla et al.,

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 moderate 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 identification of several “sub-types” of TTGs (Halla et al., 2009; Moyen, 2011a): the main differences reside in the HREE, Sr, and Nb–Ta contents (Fig. 10). Moyen (2011a) demonstrated that TTGs can be classified in three groups: the “high pressure” group has low HREE, low Nb and Ta, high Sr symptomatic of melts in equilibrium with large amounts of garnet, some rutile, but no plagioclase. In contrast, the “low pressure” group shows high HREE, Nb and Ta but lower Sr, consistent with plagioclase being stable in the residue of melting, but lesser amounts of garnet and no rutile. The “medium pressure” group is intermediate between the previous two. As plagioclase, rutile and garnet stabilities are strongly pressure-dependent; the three groups can be interpreted in terms of depth of melting. Therefore, both the experimental and the geochemical evidence indicate that the wide group of sodic Archaean granitoids (TTG s.l.) formed by melting over a large range of pressures. No single model (i.e. eclogite vs. amphibolite residue, or shallow vs. deep melting) can account for the whole group, nor is applicable as a unique model for the generation of the Archaean continental crust. Rather, it appears that diverse situations, probably corresponding to contrasting tectonic scenarios, did exist in the Archaean.

200 400 600 800 1000 1200

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Fig. 12. Different types of TTGs, identified by their geochemical features (Moyen, 2011a). The dark blue, light blue and green dots correspond respectively to high, medium and low pressure TTGs.

J.-F. Moyen, H. Martin / Lithos 148 (2012) 312–336

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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 tholeiitic amphibolite. When residual garnet is absent (green symbols) magma HREE contents remain similar to source; when garnet is present (dark blue symbols), magma becomes strongly depleted in HREE, resulting in TTG-like patterns (FSS is from Rapp et al., 1991 and D15 from Wolf and Wyllie, 1994).

(Fig. 13) that significant HREE depletions and REE fractionation in liquids resulting from metabasalt melting do require garnet to be stable in sufficient amounts in the residue.

5. The 1990s to 2000s: link with adakites Adakites were recognised in the early 1990s as a significant type of modern arc lavas (Defant and Drummond, 1990; Drummond and Defant, 1990), that show sufficient similarities to Archaean TTGs to be regarded as possible modern equivalents. Therefore, unravelling the petrogenesis of adakites can bring information on the origin of Archaean TTGs (Martin, 1999).

5.1. Adakite definition and the adakite–TTG comparison

The “conclusion” that adakites and TTGs formed under similar conditions has formed a strong basis for interpretations of the tectonic 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 subduction zone magmatism (irrespective of the details of the petrogenetic processes operating), and therefore it is tempting to conclude that TTGs are also subduction-related. Unfortunately, this logic has also been seriously misused, mainly because a wide range of igneous rock has been incorrectly described as “adakite” in the literature (Fig. 14; also see (Moyen, 2009)). This leads to spurious reasoning, whereby a rock, incorrectly classified as “adakite” is then demonstrated to be unrelated to subduction (Coldwell, 2008; Ding et al., 2011; Guo et al., 2007; Hastie et al., 2010; Li et al., 2011; Xu et al., 2006), leading to the conclusion that “adakites are not related to subduction processes” and therefore that “TTGs are not subduction-related either”. In the

Adakites, as originally defined (Defant and Drummond, 1990) are

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0.5

K2O/Na2O

1.5

High-silica adakites Low-silica adakites Archaean «adakites» Other «adakites»

0.0

This description can be compared to Martin's (1994) definition for the TTG suite, presented in Section 2.2. Obviously, the two definitions are very similar and the key points that distinguish TTG and adakites from common granitoids or arc lavas respectively are the same, i.e. the intermediate to acid composition, with the lack of mafic components; 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 definitions of both adakite and TTG is the “arc” signature of either rock group, with a typical LILE/HFSE decoupling resulting in Nb–Ta negative anomalies, low Nb/Th ratios, etc.; and a calc-alkaline nature. It is worth pointing out that every aspect of the description above is an integral part of the definition of adakite (or TTG), and that no rock should be called adakite (or TTG) if it does not match the whole definition. There is, for instance, a disturbing trend to call “any rock with elevated Sr/Y or La/Yb ratios adakite” or “adakitic”, regardless of the other geochemical characteristics (see discussion in Moyen, 2009). Bearing this in mind, there is a clear, first order similarity between TTG and adakites (Martin, 1999; Martin et al., 2005), suggesting that both rock types formed by very similar processes.

2.0

“Andesites, dacites and sodic rhyolites (dacites being most common products), or their intrusive equivalents (tonalites and trondhjemites) (…) characterized by >56% SiO2, >15% Al2O3 (rarely lower), usually b3% MgO (rarely above 6% MgO), low Y and HREE relative to islandarc andesites, dacites and rhyolites (for example, Y and Yb b18 and 1.9 ppm respectively), high Sr relative to island-arc ADR (=andesites, dacites and rhyolites) (typically >400 ppm), low high-field strength elements (HFSEs) as in most island arc ADR”.

50

55

60

65

70

75

80

SiO2 Fig. 14. Various types of “Adakites” (modified from Moyen, 2009). The different symbols correspond to different rocks described as “adakite” or “adakitic” in the original publications. The yellow field shows “ordinary” arc rocks. The dark purple data points are the “high silica adakites” (HSA) of Martin et al. (2005), i.e. the lavas that do actually match the original definition of adakites; light purple points are “low silica adakites” (LSA).

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ROCK / PRIMITIVE MANTLE

324

100

10

1

0.1

LSA HSA TTG

0.01 Rb Ba Th U K Nb Ta La Ce Sr Nd Zr Hf Sm Ti Gd Dy Y Er Yb Sc V Cr Ni 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 are significantly different. For instance, LSA display a positive Sr anomaly that does not exist in TTG and HSA, similarly LSA have higher concentrations of REE and transition element than TTG and HSA.

following discussion, we will, therefore, ignore all the “adakites” that do not correspond to the proper definition. 5.2. Are adakites and TTG really similar? Following the initial discovery, in the 1990s, that adakites are convincingly 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 the two rock types. In fact, based on geochemical considerations, Martin et al. (2005) discriminated two groups of adakites: “high-SiO2 adakites” (HSA) assumed to be generated by partial melting of a hydrous basalt and “low-SiO2 adakites” (LSA) considered as partial melts of a peridotite metasomatised by HSA. The two groups of adakites differ by several characteristics, one of them being their Mg# (61 in LSA and 48 In HAS). The two groups also differ by their Ni and 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 respectively) are very similar to HSA, but strongly differ from LSA. Fig. 15 also shows that Sr content is the same in both TTG and HSA (where this element does not correspond to any positive or negative anomaly), while LSA display a strong positive anomaly, matching much higher contents in LSA (Sraverage = 2051 ppm) compared with both HSA and TTG (565 and 400–500 ppm respectively). Finally, the LREE content of LSA is greater than in both HSA and TTG. Conclusively, it appears that most of the differences reported between 8

MgO %

a

adakites and TTG come from the grouping of two distinct rock types, LSA and HSA, which are obviously different. There are, on the other hand, 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 to have lower MgO (and Cr, Ni) than HSAs or TTGs at comparable SiO2 contents (Fig. 16). This feature is classically interpreted as reflecting interactions between the primary adakite/TTG melt and the mantle wedge (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 than that of the surrounding rocks, such that it tends to ascend towards the 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 subduction zone environments where oceanic lithosphere subducts under mantle wedge (see Section 7 for further discussion on possible tectonic scenarios). Conversely, the lack of clear geochemical evidence for melt–mantle interactions (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 mantle slice, thereby arguing against a “normal” subduction. Proposed alternatives include melting of the lower portion of a thick mafic crust (Atherton and Petford, 1993; Bédard et al., 2003; Condie, 2005b;

MgO %

b

7

(3)

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T < 3 Ga T > 3.0 Ga T < 3.5 Ga T > 3.5 Ga

5 4 3 2

(2)

(4)

(A)

Experimental liquids

1

(1)

(B)

0 40

50

60

70

SiO2%

SiO2%

Fig. 16. TTG and experimental melts: MgO vs. SiO2 systematics. a) Comparison between experimental melts and TTGs (grouped by age: Martin and Moyen, 2002) showing the generally more magnesian nature of TTGs compared to experiments. b) How mantle–melt interactions would appear in Harker type diagrams, SiO2 vs. MgO. Trend (A) is a differentiation trend (irrespective of the actual process generating it). Sample (1) has higher MgO than sample (2), but this does not reflect any sort of mantle interactions. Rather, it simply corresponds to a different position on the trend. Interactions of magma (2) with the mantle will shift it to position (3); from this point onwards, the magma will define a new trend (B). Sample (4) on this new trend actually has lower MgO than sample (2). Yet samples (3) and (4) have experienced interactions with the mantle, not sample (2)! This demonstrates that 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 a sample belongs). If the two yellow and green fields represent two sample sets, it is not immediately intuitive that the mantle-contaminated set is group (B).

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Hastie et al., 2010; Whalen et al., 2002); or a “flat subduction” scenario (Smithies, 2000; Smithies et al., 2003), with the subducted crust creeping under the upper plate without a proper mantle wedge (Gutscher et al., 2000). It is, therefore, critical to understand the physical and petrological mechanisms of mantle–melt interactions, as well as their possible geochemical consequences in order to decide whether (some?) TTG suites 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 interactions. Firstly, several groups (Condie, 2005b; Martin and Moyen, 2002; Smithies, 2000) described a secular evolution pattern, with Palaeoarchaean TTGs having lower Mg#, Ni and Cr than their Neoarchaean counterparts. They interpret this evolution in terms of progressively deeper basalt melting during the Archaean due to the progressive cooling of the Earth. Indeed, a colder Earth would result in lower geothermal gradients, such that slab melting would take place at greater depth and consequently under a thicker mantle peridotite slice, resulting in more efficient TTG melt–mantle peridotite interactions. Since the middle of 1980s, plutons referred as “phenocryst-bearing” or as “Mg-rich granodiorites” were identified amongst the late Archaean intrusions. Because the major element geochemistry of these rocks resembled that of Miocene high-Mg Andesite (Sanukite) from the Setouchi 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 Archaean craton (Bédard, 1996; Champion and Smithies, 1999; Halla, 2005; Heilimo et al., 2010; Laurent et al., 2011; Leite et al., 2004; LobachZhuchenko et al., 2005, 2008; Martin et al., 2010b; Moyen et al., 2001, 2003; Samsonov et al., 2005; Smithies and Champion, 2000; Stern and Hanson, 1991). Sanukitoids are monzodiorites to granodiorites, with an incompatible element signature similar to that of TTGs, but they are more potassic and have higher compatible elements contents (for a given SiO2 content). Sanukitoids are similar to the “low silica adakites” (Martin et al., 2005). Based on these chemical features, a common model for the origin of sanukitoids involves interactions between TTG melts and the mantle (be it in the form of contamination of the magma during ascent, or of metasomatism and subsequent melting of the peridotite, Martin et al., 2010b), i.e. a process similar to the melt– mantle interactions discussed here, with more pronounced and efficient interactions. Lastly, there is also physical evidence for primary adakite/TTG melt and peridotite interactions. Adakitic (HSA) glassy inclusions have been described in olivine crystals from ultramafic mantle xenoliths in lavas from the Batan Islands, Philippines (Schiano et al., 1995). Interstitial adakitic glass is found in mantle xenoliths from the southern Andes (Kilian and Stern, 2002), and MgO and Cr-enriched adakitic veins cut across mantle xenoliths in Kamchatka volcanoes (Kepezhinskas et al., 1995, 1996). 6.2. Petrological studies of TTG/mantle interactions and effects on the melt compositions Most of the evidence presented above is indirect, and we have only limited understanding of the petrological or physical processes of melt–peridotite interactions, or of their effects on the melt's composition. Firstly, the melt transfer mechanisms across the mantle are poorly understood. Interactions would occur assuming that the melt percolates and permeates the mantle at the grain scale, therefore allowing for important interactions. On the contrary, should melt ascend by fracturing and dyking, the scope of chemical interactions with the country rocks becomes very limited. Prouteau et al. (2001) also proposed that in this case the magma conduit will be armoured with

325

metasomatised peridotite, which will then become non-reactive. Studies of magma transfer in active margins suggests (i) a short transfer 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 zones in narrow regions with high magma fluxes (Grégoire et al., 2006; Rabinowicz et al., 1987, 2001). Both lines of evidence suggest that the opportunities for melt–mantle interactions may not be widespread, even in a subduction scenario. Secondly, the petrologic (and geochemical) effects of such interactions on the melt are poorly known. While conventional logic suggests that the net effect might be an increase in compatible element (Mg, Cr, Ni) contents and related ratios (Mg#) in the melt, the amount 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 the melt–mantle reactions (e.g. Hoffer et al., 2008; Prouteau et al., 2001; Rapp et al., 1999, 2006, 2010; Sen and Dunn, 1994; Yaxley and Green, 1998) and their effect on the melt's composition. Rapp et al. (1999, 2010) studied experimentally the origin of sanukitoids. In their experiments, the starting material consisted of TTG glass and crushed peridotite. The TTG reacted with the peridotite by incongruent reactions forming orthopyroxene and amphibole, together with a melt enriched in compatible elements but relatively unaltered trace elements patterns and ratios. Geochemical modelling of such interactions (Moyen, 2009) leads to the same conclusion. Although this study focussed on sanukitoids, it can be expected that less pronounced 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 interactions comes from high Mg, Ni and Cr contents in TTGs, for which there may be alternative explanations. For example, higher degrees of melting of the same source will give more mafic melts. Likewise, mixing with mafic magmas will have the same result. Entrainment of peritectic minerals (Stevens et al., 2007) will also yield a similar increase in compatible elements (compared to pristine melts). Lastly, high contents should not be confused with elevated ratios (for instance high Mg#). A high Mg# can be achieved by increasing Mg, but also by decreasing Fe contents. Relatively leucocratic, high-Mg# melts can be generated by high pressure melting (with Fe-rich omphacitic clinopyroxene as a residual phase (Laurie and Stevens, 2012; Moyen, 2011a)); or by melt segregation and interactions with the continental crust during ascent (Getsinger et al., 2009). Mg# does therefore not carry the same information than absolute MgO contents, and specifically is no clear marker of the role of a mafic/ultramafic component. Finally, it is worth stressing that making petrogenetic interpretations based upon only a few elements is always hazardous. Indeed, MgO, Cr and Ni are very strongly negatively correlated to SiO2 in any magmatic series, simply reflecting chemical evolution during differentiation. Therefore, comparing the MgO (for instance) content of rocks at different SiO2 levels is uninformative and misleading. 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 of the degree of differentiation (largely correlated to temperature, be it during melting or fractionation). Melt–mantle interactions would be evidenced by samples “shifting” trend, i.e. moving to a more magnesian line of evolution. The absolute values on the other hand would cover the same range in both cases. This demonstrates that 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 a sample belongs).

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7. Geodynamic implications An underlying question, present in all of the debate on TTG petrogenesis, is in what kind of geodynamic environment did these rocks form? From a pure geochemical point of view, TTGs are similar to arc rocks (although TTGs are notably depleted in HREEs and have no negative Sr anomaly compared to typical arc lavas; Fig. 5b), which has been widely used as evidence for a subduction-related origin. However, the basic information provided by TTG geochemistry is simply that these magmatic rocks derive from melting of a hydrous mafic source. In addition, the depth of melting must be great enough to stabilise garnet and sequester HREE — i.e., pressure must be >12– 15 kbar (Moyen and Stevens, 2006). Consequently, any environment allowing hydrous mafic rocks to melt at high pressure would be equally suitable to form “TTG-like” magmas, for instance an oceanic plateau (Willbold et al., 2009) or even a mid-oceanic ridge (Rollinson, 2009), if the mafic crust is thick enough. This purely comparative approach is therefore a dead-end, and a deeper understanding of TTG genesis must be sought. However, there is more than a simple qualitative link between the pressure 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 “subgroups”, one of which does require large amounts of garnet, together with rutile, in the residue. In fact the same geochemical effect can be achieved if garnet and rutile are phases fractionating during cooling and crystallisation (Macpherson et al., 2006), but the pressure requirements are the same for both models. In turn, this implies an evolution at pressures of about 20 kbar. This “high pressure TTG” group (or low HREE group in Halla et al., 2009) therefore requires the burial of 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 debated, both on theoretical grounds and geological observation. The primary cause for potentially different geodynamical behaviour in the Archaean is the change in thermal regime (Brown, 2006; van Hunen and van den Berg, 2008). The Archaean Earth produced two to three times the amount of today's radiogenic heat, and despite significant uncertainties in its early thermal history (e.g. Herzberg et al., 2010; Korenaga, 2006), the Archaean mantle was most likely hotter than it is today, as evidenced by liquidus temperatures and MgO contents of basaltic lavas through time (Abbott et al., 1994; Grove and Parman, 2004; Herzberg et al., 2010; Jaupart et al., 2007), which suggests a secular cooling rate of about 100 K/Gyr. This translates into (i) more buoyant and more rapid and/or thicker oceanic crust (van Thienen et al., 2004), partly compensated by eclogitisation of buried basalt at depth (Korenaga, 2006; van Hunen and van den Berg, 2008) as well as by the possible presence of komatiitic lavas associated with oceanic basalts in the Archaean oceanic crust (Barbey and Martin, 1987); (ii) weaker lithosphere (as the material strength is strongly temperature dependent); (iii) less efficient coupling between the convective mantle and the overlying plates (O'Neill et al., 2007a, 2007b). The net result on the convective style of Earth, and therefore on the geodynamic style of the planet, is to weaken the efficiency of subduction. Possibly, the system may transition to a “stagnant lid” regime (Moresi and Solomatov, 1998), where the lithosphere (the conductive boundary layer of the system) becomes separated from the convective mantle and “floats” on top of it without being recycled in subduction zones. Between plate tectonics and stagnant lid regimes, numerical models propose that the dominant tectonic style could have been intermittent (O'Neill et al., 2007a, 2007b), with the weak slab 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 lithospheric plates. Thus the subduction dynamics would be different (Di Giuseppe et al., 2008; Sizova et al., 2010) with less coherent slabs (van Hunen and van den Berg, 2008), smaller scale and discontinuous subduction. It is clear that if there was subduction during the Archaean, its main features were significantly different, such that many of the geological markers used in recent terrains may well become meaningless. Consequently, two main views are possible regarding Archaean tectonics. The first interpretation is that the Archaean tectonic style did not involve any sort of plate tectonics — i.e., the thermal boundary layer of the convecting system was not broken into plates integrated into a mantle-scale convection system, through subduction zones and mid-oceanic spreading centres. Alternately, one may propose that some kind of plate tectonics operated — i.e., the thermal boundary layer was mobile and recycled into the mantle along plate boundaries, but that the shape, the size, the thermal conditions or/and the timing of subductions were significantly different from the modern Earth. In addition, the Archaean aeon lasted 1.5 Ga and during this period, Earth cooled substantially, such that one could expect some change in “plate tectonics” modalities through Archaean times. Geological evidences for, or against, Archaean plate tectonics have been discussed in so many papers that a detailed account is beyond the scope of this work. The existence of Archaean subduction zones (and, therefore, plate tectonics) is supported primarily by the existence of a range of igneous rocks with an arc or arc-like composition including boninites (Smithies et al., 2004), shoshonites (Kerrich and Ludden, 2000), Nb-enriched basalts, and magnesian andesites and adakites (Polat and Kerrich, 2004). However it must be noted that the only requirement to form these rocks is the burial and dehydration of surface rocks under a mantle wedge, irrespective of the actual size and geometry of the buried portions of rock. Additional evidence includes the map pattern of many Archaean provinces (Czarnota et al., 2010; Percival et al., 2002; Poujol et al., 2003), reminiscent of lateral accretion of stitched terranes (Coney et al., 1980), as well as the existence of flat seismic reflectors, possibly representing fossil subduction planes (de Wit and Tinker, 2004; Goleby et al., 2004; Ludden and Hynes, 2000). Finally, the existence of Archaean highpressure and medium to low temperature metamorphism (Block et al., 2012; Moyen et al., 2006; Saha et al., 2010, 2011; Volodichev et al., 2004), sometimes associated with “hotter” metamorphic rocks (Moyen et al., 2006; Stevens and Moyen, 2007) in Archaean analogues of paired metamorphic belts (Banno and Nakajima, 1992; Brown, 2002; Patrick and Day, 1995), suggests the existence of some form of subduction process in the Archaean. In contrast, key indicators for modern plate tectonic processes (and in particular, subduction) are missing (Bédard et al., 2012; Hamilton, 1998, 2003; Stern, 2005). No undisputed Archaean ophiolites are known (Stern, 2007); andesites, blueschists and (in-situ) eclogites are also 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 bulk coaxial shortening of the crust (Chardon et al., 2009) and gravitydriven tectonics involving the sinking of dense greenstone belts into the soft basement (Bouhallier et al., 1995; Collins et al., 1998; Gorman et al., 1978), both suggesting deformation of a hot, “soft” Archaean lithosphere (Choukroune et al., 1995) rather than plate boundary processes. 7.2. TTGs as “non plate” magmas Assuming the absence of plate tectonics in the Archaean, the petrological requirements to form TTGs can be accommodated in several ways. Indeed, the lack of plate tectonics does not imply a lack of mantle convection; nor does it mean that mantle convection did not affect

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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 of the convecting system. However, mantle convection is able to affect the lithospheric “lid” in at least two ways (Fig. 17). Firstly, upwelling zones would cause mantle melting (in a scenario similar to modern plumes and oceanic plateaus), resulting in the accumulation of large volumes of mafic magmas forming potentially thick piles. Although the low viscosity of the lower crust will place a limit to the maximum thickness of these piles (England and Bickle, 1984; Rey and Houseman, 2006), it is possible to accumulate some 35–40 km of basalts, and therefore to reach the thickness required for the formation of at least the low-pressure TTGs (TTG s.l., high HREE in Table 2). Melting of the base of the plateau crust (heated by 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” related mafic to ultramafic lavas, similar to the situation observed in some Archaean cratons (and sometimes described as “plume–arc interactions”, Polat and Kerrich, 2001). This is neatly illustrated by the example of Kroksfjordur dacites in Iceland (Willbold et al., 2009). There, due to the interaction between the mid Atlantic ridge and a mantle plume, the crust can be 40 km thick (Kaban et al., 2002). However, despite their general similarity with TTGs s.l., the generated felsic magmas 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 account for the genesis of “high pressure” TTGs (s.s.), requiring a much greater depth of melting. Finally, basalts at the base of a thick plateau-type pile are unlikely to become hydrated; therefore, the resulting melts will, in general, not be TTG (or even I-type granites). This is illustrated by the fact that the typical felsic magma in any oceanic plateau is syenitic rather than I-type (Cousens et al., 2003; Gagnevin et al., 2003; Giret, 1990; Marsh et al., 1991; Martin et al., 2008; Shamberger and Hammer, 2006). Consequently, melting of a thick basaltic plateau can 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 also occur, either above downwelling parts of the mantle convective system (Kröner and Layer, 1992); or as increasingly dense lower crust detaches and sinks into the mantle. Increasing density can be a feature of granulite-facies conditions; or, more efficiently, a result of the formation of an eclogite residue through a first stage melting of the lower crust (Bédard, 2006). In any case, fragments of mafic rocks would sink into the mantle, where they can devolatilise (if they still have some water left), melt, and interact with the mantle in exactly the same way that subduction operates. Indeed, from a

Z R

10 Fig. 17. Non-subduction models that fit the petrological constrains. a) Over a downwelling zone, portions of the lower crust (possibly thickened by tectonic stacking and/or plume activity) delaminate and sink into the mantle. Mafic rocks are therefore heated and carried into the garnet stability field where they can melt. b) An upwelling mantle plume generates a thick oceanic plateau crust the base of which can melt to generate 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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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, grey band corresponding to the range of published estimates) as well as hydrous phases of the hydrated oceanic crust (A: anthophyllite; C: chlorite; Ta: Talc; Z: zoisite; Tr: tremolite); the “wet” and “dry” solidus for basalts; and the typical along-slab geotherms in both the modern and Archaean situation. The yellow field is the likely field for TTG genesis. b–d: cartoons corresponding to an early Archaean (b), late Archaean (c) and modern (d) situation. During the Archaean, steepening geotherms result in progressively deeper melting of the slab; in the modern Earth, the slab cannot melt (panel a) but dehydrates and yields common arc suites.. Same colour code and abbreviations as Fig. 17.

petrological point of view, the processes operating are exactly the same, whether in a true, large slab connected to an oceanic bottom; or in a small fragment of mafic rocks isolated in the mantle. Differences could be found in terms of the size of the mafic body (and therefore the rate of heating, the volume of magmas generated, etc.) but these aspects are not visible to petrology or geochemistry. Delamination 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 model The tectonic model that dominates the literature on TTG petrogenesis, however, is probably the “hot subduction” model (Martin, 1986; Fig. 18). Historically, this model is based on (i) the “arc” signature of TTGs. However, this is a weak argument, as it is possible to generate a “near TTG”, arc signature away from any arc (Rollinson, 2009; Willbold et al., 2009). (ii) the requirement for melting a hydrous basalt in the garnet stability field, i.e. >12 kbar. While more convincing than the previous argument this does not rule out the possibility of melting at the base of the crust that could conceivably reach a thickness of 35–40 km. (iii) the need for pressures of ca. 20 kbar to form high pressure TTGs. Such pressures are extremely unlikely to be achieved within the crust, especially in an Archaean crust for which the higher heat flux from the mantle would result in a weaker crust unable to support a great thickness (England and Bickle, 1984; Rey and 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 a requirement; (iv) the necessity to have huge volumes of hydrous basalts at mantle depth; (v) the fact that present day HSA, essentially similar to TTGs as discussed above, are found only in modern subduction zones; and (vi) the existence of interactions between TTG magmas (or “slab melts”) and mantle peridotite, likely in TTGs and

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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 probably valid petrogenetic processes for various rocks within the TTG s.l. group. Martin and Moyen (2002) proposed that the depth of hydrous basalt melting changed in the course of Archaean times and Moyen (2011a) showed that the resulting TTGs (s.l; sodic granitoids) can be subdivided into three sub-groups (see Section 3.3 and Fig. 12), corresponding to different depths of melting. Since the melting temperature is relatively constant regardless of pressure, this translates into melting along very different geothermal gradients (Fig. 19), from ca. 10–12 °C/km for the high-pressure group up to >30 °C/km for the low-pressure group. Regardless of the scenario preferred for Archaean tectonics, this is difficult to reconcile with a unique geodynamic setting, and it is therefore likely that TTGs s.l. can form in a variety of environments. In Moyen's (2011) database, the high-pressure group (the most likely to be subduction-related) represents some 20% of the sodic granitoids (corresponding to 10% of the grey gneisses); the lowpressure group (probably plateau-related; typically with HREE contents too high to match the TTG definition used in this work) corresponds 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 a plateau situation but probably too hot for subduction, even taking into account the fact that Archaean subductions were likely hotter. In fact, they form along a geotherm similar to the one observed in the Barberton region of south Africa at ca. 3.21 Ga (Nédélec et al., in

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7.3.2. Archaean vs. modern subduction systems Clearly, modern subduction processes do not generate large volumes of TTGs or adakites. However, both geological and geodynamic considerations indicate that the differences between Archaean and modern subduction processes are significant, such that perfect similarity should not be expected. In the subducting plate, higher mantle temperatures would favour slab melting over slab dehydration, thereby generating less andesites (related to slab dehydration), but more TTG/adakites (slab melting) (Defant and Drummond, 1990; Martin, 1986). A possible analogue can be observed in certain arc segments (such as the Austral Volcanic Zone in Southern Chile), where the subducting slab is young (or where the ridge itself is subducted), the magmas formed are adakites, rather than more typical calcalkaline suites (Bourgois et al., 1996; Guivel et al., 1999; Martin, 1999). In addition, common slab breakoff caused by thermal weakening of the subducting lithosphere (van Hunen and Moyen, 2012; van Hunen and van den Berg, 2008) would mean that the buried lithosphere was foundering and crumbling into the mantle rather than subducting as long, continuous and persistent slab. This geometry would result in short, discontinuous “bursts” of arc magmatism as opposed to continuous arc activity, as now. Short-lived events of arc volcanic and plutonic activity are indeed observed in the Archaean rock record (Moyen and van Hunen, 2012). In this view, the difference between intermittent subduction and delamination would be minor, and largely transparent to petrology and geochemistry. A possibly significant petrological difference is the degree of hydration of the basaltic rocks in both scenarios; subduction of an oceanic floor will bury hydrated basalts, whereas lower crustal delamination will affect essentially dry rocks. Consequently, the temperature of melting, melt amount and residual assemblages produced (presence or absence of amphibole) may also be expected to differ in both scenarios, and could conceivably be tracked by investigating minute geochemical or petrological differences.

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Fig. 19. P–T diagram indicating the melting conditions of the three TTG groups of Moyen (2011a) and the corresponding geotherms. The diagram shows linear geotherms for reference, as well as the range of geotherms modelled for subduction zones, in purple; the solidus of amphibolites and multivariant amphibole (or clinopyroxene + epidote) breakdown reactions (red, shaded) together with the stability limits of garnet, plagioclase and rutile (grey field, range of published curves) in mafic compositions. Although melting occurs over a large range of temperatures (700 to 1000 °C), this is still a relatively restricted range at the scale of this diagram, such that melting at different pressures as discussed previously implies melting on strikingly different geotherms (discussion in text).

press), during the exhumation and melting of high-pressure metamorphic rocks (Lana et al., 2010; Moyen et al., 2006). This suggests that the collapse of thickened orogenic (?) crust might be an option to explore for the genesis of at least some portions of the TTG s.l. group. Indeed, Miocene adakite or adakitic rocks have been described in Tibet, and related to melting of mafic lower crust in the thickened continental lithosphere (Chung et al., 2003). 8. TTG through time TTGs were originally identified as being a typically Archaean group of rocks. This contributed to the description of the Archaean as a unique period in Earth history, with distinctive geological features. However, the growing interest in TTGs; the quest for modern analogues of Archaean situations; the discovery of Hadean rocks and minerals; led to the progressive description of TTG (or TTGlike) 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 outcrop in the Northern Territories of Canada. They mainly consist of banded tonalite and granodiorite whose zircons gave ages of 4.031 ± 0.003 Ga (Bowring and Williams, 1999; Bowring et al., 1989). The major and trace element composition of the Acasta gneisses is that of TTG s.s.. They have low K2O/Na2O and fractionated REE patterns with low Yb content (YbN = 3.6) (Bowring et al., 1990). The 4.0 Ga-old tonalitic dykes cutting across the Ujaraaluk unit (previously known as Faux amphibolites, 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 were dated 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 continental crust already included TTG. To date, the Hadean record is only accessible through zircon crystals. Most of these were discovered in Western Australia at Jack Hills and Mont Narryer and are dated between 4.3 and 4.0 Ga (Cavosie et al., 2004; Compston and Pidgeon, 1986; Froude et al., 1983). Only one crystal, extracted from Jack Hills meta-quartzites gave an age of 4.404 ± 0.008 Ga (Wilde et al., 2001), which is the oldest age so far obtained on terrestrial material. The Jack Hills zircons contain several mineral 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 record a near-primary REE composition. The REE pattern of the magma from which type-1 zircon crystallised (Fig. 20) can be calculated using the zircon/magma partition coefficients determined by Hinton and Upton (1991). These patterns are fractionated (high La/Yb) with strong negative Eu and positive Ce anomalies, that Peck et al. (2001) ascribe to the unknown oxidation state during zircon crystallisation. They probably do not reflect accurately the composition of the parent magma. The first order feature remaining is, therefore, the fractionated nature of the REE patterns of the liquids in which the Jack Hills zircons crystallised. Wilde et al. (2001) consider that such REE patterns, together with the presence of muscovite and K-feldspar in the host magma, indicate that this latter must have been generated by recycling of older granitoids; the source of which must already have TTG-like fractionated patterns. Finally, an exhaustive isotopic Lu–Hf study (Blichert-Toft and Albarède, 2008) showed that the magma from which the Jack Hills zircons crystallised had 176 Lu/ 177Hf b 0.01. This is inconsistent with MORB, oceanic plateau basalts or arc magmas, but similar to Archaean TTG. The available evidence therefore suggests that Jack Hills zircons crystallised from a TTG-like granitoid melt (Blichert-Toft and Albarède, 2008; Guitreau et al., 2012; Nebel-Jacobsen et al., 2010). Consequently, TTG-like magmas appear to be amongst the most ancient types of felsic magmas 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 the abundance of TTG, progressively replaced by potassic I-type granitoids. This shift in the magmatic record has been attributed to the change from the archaic to modern style of plate tectonics (Martin,

1993; Stern, 2008) or to an episodic mantle overturn regime and possible shutdown of plate tectonics at the close of the Archaean (Condie et al., 2009; O'Neill et al., 2007b; Silver and Behn, 2008). The latter would 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–Birimian orogeny (Almeida et al., 2007; Baratoux et al., 2011; Conceição de Araújo Pinho et al., 2011; Delor et al., 2003; Enjolvy, 2008); as well as during the 0.8–0.6 Ga Pan‐African orogeny (Isseini, 2011). They are generally associated with huge volumes of either calc–alkaline juvenile magmas (De Souza et al., 2007) or products of intracrustal melting. 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 common with time, but never totally stopped. Today, TTGs are restricted to exceptional geodynamic environments, most often in association with adakitic volcanics. This is the case for the ca. 4 Ma–old Cabo Raper pluton, emplaced in the Taitao peninsula (Chile), in response of the Chile ridge subduction (Lagabrielle et al., 1994, 2000). This TTG s.l. pluton (Bourgois et al., 1996; Guivel et al., 1999) is associated with the adakites of the Austral 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 explore for a better understanding of TTGs. 9.1. The source of TTG magmas Although most workers propose, explicitly or not, MORB-like sources for Archaean mafic rocks, several lines of evidence suggest that enriched basalts are required. For instance the models of Martin (1987), described in Section 3.2, uses a source with a slightly fractionated REE pattern, like a modern E-MORB. An enriched mafic source is also required 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 enriched trace elements signature (Condie, 1981, 2005a; Hollings and Kerrich, 2006; Jahn et al., 1980; Martin, 2011; Moyen, 2011b; van Hunen and Moyen, 2012). The dominant source of Archaean TTGs was probably not MORBs, but rather an enriched tholeiitic basalt (Martin, 2011; Moyen, 2011b). It would be interesting to compare with more recent

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TTGs, and see whether this evolution mirrors an evolution in the composition of the dominant type of mafic rocks.

9.2. TTG with a long crustal residence? Whole rock radiogenic isotopic studies in TTGs initially turned out to be rather disappointing. Despite some heterogeneities, the typical TTG is near chondritic (εNd(T) ranging from −4 to + 2 for instance), and no correlation can be demonstrated with other major or trace element patterns. The slightly enriched elemental signature observed in 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 zircons has revived some interest for isotopic studies of TTGs (or grey gneisses, described as TTGs in most of the works cited here and one must bear in mind that several distinct rock types were thus probably included under this name). A recent paper by Guitreau et al. (2012) reports >12,000 Lu–Hf analyses on zircons extracted from grey gneisses, together with 141 Lu–Hf whole rock analyses on TTG s.l. samples. The age distribution in single zircons confirms the findings of e.g. Albarède (1998), Condie (1998) and Rino et al. (2004). The clustering of zircon ages suggests that crustal growth was episodic and occurred in three or four main episodes during the Archaean. Each episode of crustal growth starts with magmas having a chondritic or positive εHf progressively, the zircon's isotopic signature evolves towards negative εHf values. This evolution is interpreted in term of recycling of juvenile crust formed at the start of each “super-event”. The 176Lu/177Hf ratio of juvenile granitoids formed during each super-event did not vary throughout the Archaean, and remains close to the chondritic value of 0.0336 (Bouvier et al., 2008). Consequently, it appears that since ~3.8 Ga the extraction of continental crust has not depleted its mantle source. A possible interpretation is that TTGs are generated from a mafic precursor ultimately extracted from the undepleted (lower) mantle (a possibility also supported by the trace elements data discussed in Section 9.1), such as basalts from oceanic plateaus formed above a deep-seated mantle plume. The remaining plume residue would give rise to the depleted upper mantle. The source of granitoids formed between the main events of crustal growth has a long crustal residence time, possibly separated from the mantle several hundreds of Myr before the formation of the specific rock specimen (Guitreau et al., 2010; Kröner et al., 2011; Zeh et al., 2008, 2009). Furthermore, zircons extracted from gneisses from a specific region or terrane tend to follow a distinctive εHf vs. time evolution (Zeh et al., 2009), suggesting that all gneisses from one area, regardless of age, ultimately derive from a similar mafic precursor, extracted roughly at the same time from the mantle. Several interpretations can account for this observation: (i) the observed evolution could be dominantly a feature of recycling of older gneisses; in other words, the grey gneisses samples are actually dominated by rocks that are not TTGs but rather potassic granitoids formed by successive melting events of the earliest TTG gneisses. TTG would then form only during the discreet, juvenile super-events. This corresponds to the evolution described in the East Pilbara (Champion and Smithies, 2007), where successive granitic generations between ca. 3.5 and 2.9 Ga become more potassic, but keep a similar Nd model age, suggesting that all the granitoids emplaced during ca. 600 Myr formed by successive melting and refining of the earliest continental nucleus, made of juvenile TTGs intruded between 3.5 and 3.4 Ga. (ii) Assuming that the sampled grey gneisses were actually proper TTGs and were generated by melting of mafic rocks, this would imply that the mafic precursor was extracted from the mantle long before it melted to yield TTGs. Different tectonic scenarios can allow such a model; for instance, extraction of thick mafic plateaus, that survive for a long period before being subducted (Guitreau et al., 2012; Martin, 2011; Moyen, 2011b) and melted to form TTGs

or a “stagnant lid” geology, where a persistent basaltic shell occasionally melted to form TTGs (Fig. 17). In any case, it is difficult to see how the problem could be solved without coupling all approaches and trying to reconstruct the nature of the source and its crustal residence time as well as the conditions of melting of the studied gneiss samples. Once again, it must be stressed that naming a sample “TTG” is not enough to replace a careful petrogenetic study. 9.3. Petrology Nearly all the discussion presented here, and indeed the best part of the literature, is based on geochemical considerations. Yet, TTGs are granitoids (or meta-granitoids), and at least in some examples it would be possible to use petrological information to constrain their evolution. In particular, a salient feature of many TTG intrusions is the presence of magmatic epidote (Bédard, 2003; Moyen et al., 2007). In granitic 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 petrogenetic and tectonic models. Recent advances in granite petrology emphasised the role of the melting reactions and their products, both melt and peritectic minerals (Clemens et al., 2009, 2011; Stevens et al., 2007; Villaros et al., 2009b) in shaping the geochemistry of granitic intrusions. In highly peraluminous granites (=S-type granites s.s., Villaseca et al., 1998) A/CNK is positively correlated with Mg + Fe. This was interpreted by Stevens et al. (2007) as reflecting the fact that S-type granites are a mixture between melt and entrained peritectic garnet. In contrast, moderately or low peraluminous granites (Villaseca et al., 1998) as well as metaluminous “I-type” granites (Clemens et al., 2011) show a negative correlation between A/CNK and Fe + Mg. Clemens et al. (2011) interpreted this correlation as evidence for the entrainment of peritectic clinopyroxene together with the melt. Finally, leucogranites (felsic peraluminous in Villaseca et al., 1998) show no correlation; their compositions are reasonably similar to these of true experimental melts (Stevens et al., 2007). TTGs show either a 1.50

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negative correlation between A/CNK and Fe + Mg (Fig. 21) or for the leucocratic units (typically low HREE “high pressure” TTGs), no correlation. This is in good agreement with the low HREE content of TTGs compared to S-types: although S-type melts form in equilibrium with garnet, the resulting granites are not HREE-depleted because peritectic garnet is formed together with the melt and entrained with it. In contrast, in TTGs garnet is not entrained, perhaps because it existed as a porphyroblast in the pre-melting metamorphic rocks, and was therefore not available for entrainment. Regardless of the detailed petrogenetic processes however, it is clear that the TTG group encompasses rocks having contrasting relationships between A/CNK and Fe + Mg, and therefore having distinct petrogenesis, that should be elucidated. 9.4. Improper uses, dubious comparisons and groupings 9.4.1. Definitions, classifications and terminology Although the term of “TTG” is in wide use, there is actually no formal, official definition for this rock type. As the previous discussion has demonstrated, however, a large part of the debates surrounding the origin of TTGs is probably tied to definition issues, with different scientist using the same name for different rocks. In particular, the geology of Archaean cratonic areas is commonly described in terms of a “basement” of gneisses, intruded by plutons or overlaid by supracrustal greenstone sequences. More or less implicitly, many geologists would regard the gneissic basement as a whole, made up of “TTG” gneisses, and therefore the TTG acronym is frequently used as a synonym for “grey gneisses”. However, grey gneisses do not consist of TTGs alone, but rather include a range of granitoid (sodic and potassic) as well as miscellaneous (amphibolites, leucosomes, restites, etc.) components. In Moyen's (2011) database, a mere 50% of the samples of Archaean grey gneisses were classified as TTG. Fig. 22 demonstrates that a significant portion of the published literature on “TTGs” is actually based on rocks that should never have been classified as such.

6

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, dispute the wisdom of comparing lavas and plutonic rocks exclusively on geochemical grounds. Firstly, whenever coeval lavas and intrusive are observed (Bachmann and Bergantz, 2004) they typically do not have the same composition; suggesting that the apparent similarity might not be so significant. Secondly, lavas and intrusive rocks do not have the l

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same behaviour, in terms of petrology. Lavas are erupted, meaning that they do not cross their solidus on the ascent path, whereas plutons represent magmas trapped at depth. Therefore, it is likely that the two types of magma have some significant differences, especially in terms of water saturation (Clemens, 1998; Clemens and Droop, 1998). Thirdly, lavas (but typically not “high-SiO2 adakites”) do commonly belong to differentiation trends, spanning a rather large range of chemical compositions from basalts to intermediate or acid compositions. The same is not true for plutonic rocks (and for TTGs in particular) where the range of composition in any given pluton or intrusive suite does typically not exceed a few weight percent SiO2; demonstrating that while differentiation processes play an important role in shaping volcanic suites, they are not so significant for plutonic rocks. Finally, lavas are demonstrably pure or nearly pure liquids, with a subordinate amount of phenocrysts. The same does not apply to plutonic rocks, however, in which inheritance and transfer of solid crystals is common, including accessory minerals (inherited zircon cores), and, probably, also major minerals carrying a significant portion of the budget of elements such as Fe and Mg (Stevens et al., 2007; Villaros et al., 2009a). The geochemical signature of plutonic and volcanic rocks, even if derived from the same source, may therefore not be shaped by exactly the same processes. Again, proper petrogenetic studies are required to address this issue. 9.4.3. TTGs — tonalites, trondhjemites and granodiorites Finally, 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 unique TTG series, or does this term actually correspond to several rock types sharing similarities and typically tectonically mixed and interleaved? The evidence presented here does indeed suggest that the term “TTG” is a misnomer. There is no “TTG” series — there are tonalites (low/medium pressure melting of mafic rocks), trondhjemites (high pressure melting of mafic rocks) and granodiorites (melting of enriched crustal lithologies). 10. Conclusion Since their initial description, in the 1970s, TTGs became a matter of considerable scientific interest. The discussion on TTG origin is connected to many important topics in Earth Science, such as the differentiation of the silicate Earth, the tectonic style of the Early Earth and its secular evolution towards plate tectonics as we know it today, the origin of arc magmas in general (role of slab melting, mantle–melt interactions), the meaning of differentiation “trends” in geochemical diagrams, etc. We now have a conceptual framework in which most researchers think about TTGs. However, despite the large volume of literature available, many first order questions still remain unanswered. Large parts of the controversies are probably the result of incorrect definitions of what TTGs are, either in geochemical terms (low vs. high Al2O3 series, HREE contents, K/Na ratios, etc.), or in terms of plain field descriptions (grey gneiss complexes comprise more than TTGs, such that any study attempting to unravel TTG geochemistry based on a “grey gneiss” sample set is intrinsically flawed). After 40 years of advances in analytical techniques, with an increasing volume and diversity of data available to the researcher, it is worth noting that old fashioned field observations and petrography still hold the key to the solution of some first order problems.

Calc-alkaline

1

Acknowledgements

0

Tholeiitic

60

65

70

75

80

SiO2 Fig. 22. Grey gneisses include more than TTGs (modified from Moyen, 2011a). The dark blue, light blue and green dots are TTGs proper (as in Fig. 10). The grey symbols (see Moyen, 2011a for further details) are the other (non TTG) components of grey gneisses.

The authors are grateful to Hugh Smithies and Bor-Ming Jahn for detailed and fruitful reviews which greatly improved the quality of this manuscript. Editorial assistance by Nelson Eby is also gratefully acknowledged. This work evolved from a keynote talk given at the 23rd Congress of African Geology, held in Johannesburg in January 2011; the authors wish to thank the organizers of the conference for invitation. This review is the result of a long scientific association with TTGs and people trying to understand them, and we wish to

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