Directory UMM :Data Elmu:jurnal:P:Precambrian Research:Vol105.Issue2-4.2001:

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Protoliths of the 3.8 – 3.7 Ga Isua greenstone belt,

West Greenland

John S. Myers *

Department of Earth Sciences,Memorial Uni6ersity of Newfoundland,St.John’s,Nfld Canada,A1B3X5

Received 15 July 1999; accepted 18 October 1999

Abstract

The Isua greenstone belt (Fig. 1) contains the oldest known, relatively well preserved, metavolcanic and metasedimentary rocks on Earth. The rocks are all deformed and many were substantially altered by metasomatism, but both the deformation and metasomatism were heterogeneous. Transitional stages can be seen from relatively well preserved primary volcanic and sedimentary structures to schists in which all primary features have been obliterated. Likewise different kinds, and different episodes, of metasomatic alteration can be seen that produced a diversity of different compositions and metamorphic mineral assemblages from similar protoliths. New geological mapping has traced out gradations between the best preserved protoliths and their diverse deformed and metasomatised equivalents. By this means, the primary nature of the schists that make up most of the Isua greenstone belt was reinterpreted, and a new map that better portrays the primary nature of the rocks has been produced. The previously mapped stratigraphy was found to be of little value in understanding the geology. Stratigraphic units were defined by different and diverse criteria, such as current composition, structure, metamorphic texture, and inferred protoliths. Much of this stratigraphy represents a misinterpretation of the primary nature of the rocks. The new work indicates that most of the Isua greenstone belt consists of fault-bounded rock packages, mainly derived from basaltic and high-Mg basaltic pillow lava and pillow lava breccia, chert – BIF, and a minor component of clastic sedimentary rocks derived from chert and basaltic volcanic rocks. A previously mapped, extensive, unit of felsic volcanic rocks was found to be derived from metasomatised basaltic pillow lava and pillow breccia intruded by numerous sheets of tonalite. © 2001 Elsevier Science B.V. All rights reserved.

Keywords:Early Archaean; Tectonic evolution; Greenstone belt; Protolith interpretation; Greenland

www.elsevier.com/locate/precamres

1. Introduction

The Isua greenstone belt (also known as Isua supracrustal belt) is part of the Archaean gneiss

complex of West Greenland (Bridgwater et al., 1976). Most of this gneiss complex consists of tonalitic gneiss with minor components of granitic gneiss, layered megacrystic anorthosite complexes, amphibolite derived from basaltic volcanic rocks, metasedimentary rocks and ultramafic rocks, that all formed between 3.0 and 2.7 Ga. In the middle * Tel.: +1-709-7378417; fax:+1-709-7372589.

E-mail address:[email protected] (J.S. Myers).

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of this gneiss complex there is a belt 50 – 75 km wide, extending for 200 km northeast through Godtha˚bsfjord, that contains fragments of early Archaean rocks. Most of the latter are tonalitic gneisses (Amıˆtsoq gneiss: McGregor, 1973) that formed between 3.87 and 3.65 Ga (Nutman et al., 1996) or at ca. 3.65 Ga (Kamber and Moorbath, 1998; Whitehouse et al., 1999). These ancient

tonalitic gneisses contain fragments of

supracrustal, metavolcanic and metasedimentary rocks. Small fragments of these rocks are known as the Akilia association (McGregor and Mason, 1977) and the largest fragment is called the Isua supracrustal belt or Isua greenstone belt (Appel et al., 1998). The protolith ages of the Akilia associ-ation are controversial: as old as ca.3.87 Ga (Nut-man et al., 1996) or 3.7 – 3.65 Ga (Kamber and Moorbath, 1998; Whitehouse et al., 1999). The precise ages of the diverse components of the Isua greenstone belt are unknown, but numerous age determinations by diverse methods in several lab-oratories indicate protolith ages between 3.8 and 3.7 Ga (Moorbath et al., 1997; Nutman et al., 1997).

The Isua greenstone belt (Fig. 1) contains the best preserved, oldest known sequence of rocks that formed on the surface of the Earth. These rocks are therefore of outstanding importance in recording the oldest known terrestrial environ-ments, and provide the best opportunity for dis-covering the earliest traces of life on Earth.

The interpretation of these ancient environ-ments, as well as the search for traces of life, rely upon correct identification of the original nature of the schists that make up the greenstone belt. This in turn requires that the complex tectonic, metamorphic and metasomatic history of the rocks be unravelled in order to determine both the original nature and relationships of the compo-nents of the greenstone belt.

The volcanic and sedimentary rocks from which the greenstone belt was derived were intruded by sheets of tonalite and several generations of doler-ite dykes. The rocks were repeatedly deformed and recrystallised in upper greenschist to lower amphibolite facies conditions. All the components

of the greenstone belt are strongly deformed and most are schistose.

Most research on the Isua greenstone belt since 1984 has been described in the context of the geological map and stratigraphy of Nutman et al. (1984) and Nutman (1986). These authors divided the greenstone belt into nine formations and two kinds of intrusive rocks: ultramafic rocks and a major unit called ‘garbenschiefer amphibolite’ that was interpreted as a Mg – Al-rich basic intrusion.

Rose et al. (1996) made a detailed study of the ‘calc-silicate formation’ of Nutman (1986). In contrast to Nutman (1986) they concluded that this unit was not derived from calcareous chemi-cal sediments but by metasomatism ‘‘where fluids flowed across the contacts between ultramafic bodies and felsic or metabasaltic country rocks at deep crustal levels’’.

The whole stratigraphy of Nutman (1986) was questioned by Rosing et al. (1996) who considered that the rock sequence was derived from basalt and banded iron formation, intruded by ultra-mafic sills and sheets of tonalite – granite, and heterogeneously altered by metasomatism. In con-trast to previous interpretations that the Isua sequence formed in a shallow water, platform environment, with clastic sedimentary rocks in-cluding conglomerates, and calcareous chemical precipitates, Rosing et al. (1996) suggested that the Isua sequence could have originated in an oceanic environment. These authors also reinter-preted the ‘‘garbenschiefer amphibolite’’ as a unit of mixed volcanic and sedimentary origin rather than an intrusion as suggested by Nutman et al. (1984) and Nutman (1986, 1997).

The new work described here generally sup-ports the reinterpretation of the Isua ‘stratigra-phy’ by Rosing et al. (1996). Field evidence is presented of the tectonic and metasomatic transi-tions by which a diversity of metamorphic rocks were generated from a few, relatively uniform, protoliths. The distribution of these protoliths, and of regionally extensive zones of metasomatic alteration, are shown on tectonostratigraphic maps of two portions of the Isua greenstone belt (Figs. 2 and 3).

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2. Isua greenstone belt — previous work and stratigraphy

The Isua greenstone belt (Fig. 1) is located on the edge of the inland ice cap, 150 km northeast of Nuuk. The greenstones form an arcuate belt 35 km long that is truncated to the northwest by the Ataneq fault (McGregor, 1979). This fault is linked in the southwest to the Ivinnguit fault that was interpreted by McGregor et al. (1990) as a ca. 2.72 – 2.7 Ga terrane boundary between the Akulleq terrane, containing the Isua greenstone belt and other early Archaean rocks, and the Akia terrane to the northwest dominated by ca. 3.2 – 2.98 Ga tonalitic gneiss.

2.1. Pre6ious work

A substantial amount of research has been car-ried out on the Isua greenstone belt since the great antiquity of these rocks was first established by Moorbath et al. (1972, 1973) who obtained a

Pb/Pb whole rock age of 3760970 Ma

(subse-quently recalculated to 3710970 Ma, see

Moor-bath and Whitehouse, 1996) on the banded iron

formation, and a Rb/Sr whole rock age of 37009

140 Ma from the tonalitic gneiss. The main fea-tures of the geology were first described by Bridgwater and McGregor (1974). They com-pared the tonalitic gneiss to the Amıˆtsoq gneiss of the Godtha˚b (Nuuk) region (McGregor, 1973), and the dykes that cut both the gneiss and the

supracrustal rocks to Ameralik dykes. The

supracrustal rocks were mapped and described by Allaart (1976), and further description and inter-pretation were given by Bridgwater et al. (1976). They discussed various interpretations of the quartzo – feldspathic schists interleaved with the mafic and ultramafic rocks and concluded that they were derived from acid volcanic rocks.

There was a surge of research activity during the late 1970s to early 1980s on a variety of topics including: stratigraphy and sedimentology (Dim-roth, 1982; Nutman et al., 1984); structure (James, 1976); petrology, mineralogy and geochemistry (Schidlowski et al., 1979; Gill et al., 1981; Boak et

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al., 1983); metamorphism (Boak and Dymek, 1982); geochronology (Moorbath et al., 1975; Baadsgaard, 1976; Michard-Vitrac et al., 1977; Hamilton et al., 1978), oxygen and sulphur iso-tope studies (Oehler and Smith, 1977; Oskvarek and Perry, 1976; Perry and Ahmad, 1977) and organic chemistry (Nagy et al., 1975, 1977). The Isua belt was remapped by Nutman in 1980 – 82 at a scale of 1:10000, and the stratigraphy presented by Nutman et al. (1984) and Nutman (1986) has remained the basis for all subsequent research, and was recently reiterated with new geochronol-ogy by Nutman (1997) and Nutman et al. (1997), (1998).

2.2. Nutman stratigraphy

Nutman (1986) defined two stratigraphic se-quences (A and B). Sequence A formed ‘a coher-ent stratigraphy…along the whole length of the Isua supracrustal belt’ whereas part of sequence B was cut out by a fault (Nutman, 1986, p. 10). He defined sequence A as comprising six formations ‘in ascending order’ (Fig. 3a): (A1) amphibolite formation of mainly banded amphibolite; (A2) lower banded iron formation, characterised by magnetite iron formation; (A3) variegated schist formation consisting of ‘amphibolites, felsic rocks, metapelites and metacherts’; (A4) upper banded iron formation of ‘magnetite iron formation and metachert’; (A5) calc-silicate formation compris-ing ‘carbonates, quartzites, calc-silicate rocks and amphibolites’; and (A6) felsic formation of ‘‘pre-dominantly layered metasediments’’. The accom-panying map (Nutman, 1986, plate 1) included a seventh stratigraphic unit of ‘‘undifferentiated variegated schist formation to amphibolite forma-tion’’, placed beneath the amphibolite formation (A1). ‘Sequence B comprises in ascending order: (B1) felsic formation’ of ‘predominantly metasedi-ments’ ‘and (B2) mica schist formation’ ‘predomi-nantly Fe – Mg-rich mica schists’ (Nutman, 1986). He described several alternative explanations of the original relationship between sequences A and B.

Nutman (1986) interpreted two additional units as mafic intrusions. ‘garbenschiefer amphibolite’ comprising ‘units up to more than 1 km broad

that form approx. 25% of the Isua supracrustal belt’ and ultramafic rocks (Fig. 3a). The ‘garben-schiefer amphibolite’ was described as being

char-acterised by ‘well-developed garbenschiefer

texture of amphiboles on its foliation surfaces’, and as being ‘‘slightly discordant to lithological layering in adjacent rocks’’. This unit was inter-preted as a ‘gabbroic, possibly sill-like intru-sion(s)’. The ultramafic rocks were described as being located ‘at most stratigraphic levels in the Isua supracrustal belt’ and as being ‘slightly dis-cordant to layering of the supracrustal sequences’. They were interpreted as being derived from in-trusions of mainly dunite and peridotite.

2.3. Reappraisal of the Nutman stratigraphy Doubt was cast over the geological map and stratigraphy of Nutman (1986) by Rose et al. (1996), following a detailed study of the metacar-bonate rocks. These authors concluded that the calc-silicate formation was metasomatic in origin. Further reappraisal of the supracrustal rocks by Rosing et al. (1996) started to demolish the rest of this stratigraphic edifice. The carbonate-rich gar-net – hornblende – biotite schists that are the main component of the ‘variegated schist formation’ were found to be gradational into amphibolite, and to have similar, basaltic, rare earth element content. These schists were reinterpreted as meta-somatic alteration products of amphibolite, rather than being of sedimentary origin. Likewise grada-tions were described between leucocratic schists, mainly quartz – feldspar – muscovite – biotite – car-bonate rocks, that comprise the ‘felsic formation’, and discordant sheets of tonalitic and granitic gneiss. Rosing et al. (1996) therefore concluded that the ‘felsic formation’ was derived by metaso-matic alteration of intrusive tonalite and granite, rather than being a sequence of metamorphosed felsic volcanic and epiclastic rocks. In contrast, part of the ‘garbenschiefer amphibolite’ unit was found to contain components of sedimentary origin that passed gradually into the dominant garbenschiefer-textured magnesian schists. This unit was reinterpreted as of volcano-sedimentary, chiefly basaltic, origin and an integral part of the stratigraphy, rather than being an intrusion,

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re-cently described as metagabbro by Nutman (1997).

A consequence of this reappraisal of the stratig-raphy was that the simple isoclinal syncline pro-posed as the dominant structure of the greenstone belt by Nutman et al. (1984) and Nutman et al. (1996), Nutman (1997) was also unfounded. Ros-ing et al. (1996) considered that the dominant protoliths of the greenstone belt were basalt and banded iron formation, intruded by dunitic sills. They recognised that the sequence was deformed and metamorphosed, including being sliced up by faults, before being intruded by tonalite and gran-ite, followed by further metasomatism, deforma-tion and metamorphism. Rosing et al. (1996) concluded that ‘pervasive carbonation and K metasomatism produced a sequence of lithologies, mimicking those found in modern platform de-posits. However, the protoliths could have origi-nated in a purely oceanic environment with no sialic detrital components’.

3. Major components of the Isua greenstone belt and their protoliths

The northeast and southwest parts of the Isua greestone belt (Fig. 1) have been remapped at a scale of 1:8000 and the main features of the geology are shown in Figs. 2 and 3b. The green-stone belt and adjacent tonalitic gneisses were

repeatedly deformed and recrystallised. The

greenstones were folded into isoclinal structures that were refolded isoclinally before different seg-ments of the greenstone belt were juxtaposed in their current relative positions by ductile faults. Subsequently the whole package of greenstone slices was deformed and folds were generated on all scales with axial surfaces inclined steeply to the southeast and fold axes and associated lineations plunging moderately to the southeast (Figs. 2 and 3b).

In the northeast (Fig. 2) the greenstone belt is composed of three fault-bounded slices, infor-mally described as northwest, central and south-east tectonic domains by Appel et al. (1998). Deformed primary features (depositional struc-tures and macroscopic texstruc-tures) are widely

pre-served in the central domain, whereas in the adjacent domains most primary features have been obliterated by more intense deformation that converted most rocks to schists.

The southwest part of the greenstone belt (Fig. 3b) is also divided by ductile faults into a number of tectonic slices. Here the deformation was more heterogeneous, and although most rocks are schis-tose, a variety of primary features are locally well preserved in all the tectonic slices.

The main rock types and structures are de-scribed below:

3.1. Amphibolite deri6ed from pillow la6a and related epiclastic rocks

This is the most widespread mappable unit and is marked as ‘‘amphibolite (pillow lava)’’ on Figs. 2 and 3b. The rock mainly consists of actinolite – chlorite – talc – tremolite schist. The deformation was heterogeneous and the least deformed rocks contain pillow lava (Figs. 4 and 5) or related epiclastic structures. With increasing deformation, these rocks were converted to banded amphibolite (Figs. 6 and 7). Some deformed pillow lavas con-tain cooling collapse cavities infilled by quartz (Fig. 4) and some contain occelli (Fig. 5). In some cases the occelli are concentrated in concentric zones around the outer parts of individual pil-lows. Some of the occelli are cut by concentric cracks infilled by quartz that formed during the crystallisation of the magma.

3.2. Amphibolite (altered pillow la6a)

This widespread mappable rock type, marked

as ‘amphibolite9garnet9carbonate (altered

pil-low lava’ on Figs. 2 and 3b, mainly consists of hornblende – garnet – biotite – carbonate

(dolomite – ankerite) schist. The rocks were het-erogeneously deformed and, as in the previous unit of amphibolite, with increasing deformation all stages can be seen between amphibolite with deformed pillow lava structure and banded am-phibolite. There is a spatial association between much of this garnet – carbonate amphibolite and chert – BIF. This kind of amphibolite is located along the margins of all chert – BIF horizons and

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Fig. 2. Map showing the main features of the northeast part of the Isua greenstone belt (located on Fig. 1). Red lines indicate major faults. Dip and strike symbols relate to schistosity and transposed compositional banding, and arrows indicate the direction and plunge of fold axes and lineations.

passes gradationally away from these margins into actinolite-chlorite – talc – tremolite amphibo-lite. (The thin layers of chert – BIF in ‘amphibolite (pillow lava)’ in Fig. 2 are also bounded by gar-net-carbonate amphibolite but this is too thin to be shown on this map.)

The spatial association with chert – BIF, and the gradation in composition into actinolite –

chlor-ite – talc – tremolchlor-ite amphibolchlor-ite with similar de-formed pillow lava structures, suggests that the garnet – carbonate amphibolite represents a meta-somatically altered equivalent of the actinolite –

chlorite – talc – tremolite amphibolite. Such

alteration occurred before the last episodes of deformation and recrystallisation because both rock types are equally overprinted by these events.

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3.3. Chert-banded iron formation

Recrystallised chert and banded iron formation (BIF) form a major unit in the northeast of the

greenstone belt (Fig. 2) and thin layers within amphibolite (Fig. 3b). The BIF largely consists of alternating layers of quartz and magnetite. There are complete gradations between chert and BIF.

Fig. 3. a — Map of the southwest part of the Isua greenstone belt, simplified from Nutman (1986) (located on Fig. 1). b — New map of the southwest part of the Isua greenstone belt (located on Fig. 1). Red lines indicate major faults. Dip and strike symbols relate to schistosity and transposed compositional banding, and arrows indicate the direction and plunge of fold axes and lineations.

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Fig. 4. Deformed and recrystallised basaltic pillow lava. A deformed, quartz-filled primary collapse structure can be seen in the centre of a pillow to the right of the scale card. The pillow matrix that has been eroded away consisted of biotite and carbonate.

Where the BIF is least deformed the layers of quartz and magnetite are generally 0.2 – 1.0 cm thick and these layers reflect deformed primary layering that developed on or below the ocean floor. However even the least deformed layering is substantially modified by deformation, and most layering was folded and extended parallel to the plunge of the regionally dominant folds and asso-ciated lineations (Figs. 2, 3b and 8). In some cases individual layers of quartz were disrupted into isolated boudins or fracture-bounded tabular seg-ments during deformation that preceded the fold-ing (Fig. 8). These structures have previously been interpreted as sedimentary (‘flat pebble conglom-eratic structure’; Nutman, 1986, p. 17 – 18).

Primary conglomeratic structures occur in a chert – BIF horizon in the northeastern part of the greenstone belt, in the southern part of the central tectonic domain (Fig. 2). These conglomerates include both oligomict conglomerate with round pebbles of quartz in a matrix of mainly quartz, biotite and garnet (‘round pebble conglomeratic structure’ of Nutman, 1986, p. 18), and polymict conglomerate containing pebbles of quartz and meta-basalt in a matrix of quartz, biotite and garnet (Appel et al., 1998).

However, most chert – BIF is intensely de-formed and in the northeast part of the green-stone belt (northeast part of Fig. 2) primary layering is extensively transposed into a new tec-tonic layering (Fig. 9) and much of the chert comprises mylonite or recrystallised mylonite. In the southwest part of the greenstone belt (Fig. 3b) many thin layers of banded quartz schist and mylonite could have been derived from either primary chert or from quartz veins.

3.4. Ultramafic rocks

Ultramafic rocks are of two kinds: anthophyl-lite-rich rocks and layered serpentinites associated with metagabbro and metapyroxenite. Each kind of ultramafic rock is confined to different tectonic slices of the greenstone belt.

In the southwestern part of Fig. 3b, anthophyl-lite-rich ultramafic rocks form thin layers within actinolite – chlorite – talc – tremolite amphibolite with deformed pillow lava structures, and thicker fault-bounded layers. The apparently simple, mas-sive field appearance of these rocks results from coarse radiating clusters of anthophyllite that form most of the rocks. However this overprints a complex history of older tectonic layering and

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deformed mafic and ultramafic dykes. These ultra-mafic rocks could have originated as either ko-matiite flows or sub-volcanic intrusions.

In the northeast part of Fig. 3b, ultramafic rocks comprise serpentinite spatially associated with compositionally layered metagabbro and metapyroxenite. These ultramafic rocks appear to be derived from layered dunite – peridotite – pyrox-enite – gabbro intrusions.

3.5. Quartzo-feldspathic schist

Prominent layers of quartzo – feldspathic schist (Figs. 2 and 3) have previously been interpreted as derived from felsic volcanic rocks (Allaart, 1976; Nutman, 1986). Most of these rocks are intensely deformed schists or mylonites, but the amount of

Fig. 6. Banded amphibolite derived from strongly deformed recrystallised pillow lava with flattened and extended dark pillow cores and pale pillow rims.

Fig. 5. Deformed and recrystallised high-Mg basalt pillow lava with inter-pillow quartz. The pale spots within the pillows are deformed occelli.

deformation in the thickest layer in Fig. 3b de-creases towards the east and the schist passes gradationally into tonalitic gneiss. To the north-west in Fig. 3b, thin attenuated and boudinaged layers of quartzo – feldspathic schist and mylonite occur within garnet – biotite – carbonate amphibo-lite with discoidal bodies of quartz (Fig. 10). These rocks were previously interpreted as pyro-clastic felsic volcanic rocks (Allaart, 1976; Nut-man, 1986), but are here interpreted as deformed and recrystallised pillow lavas (with discoidal quartz representing pillow matrices) and de-formed, tectonically disrupted, sheets of recrys-tallised tonalite, derived from veins intruded into the pillow lavas, as these features can be observed in the least strained parts of this unit. In other places where the pillow lavas suffered a greater

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Fig. 7. Banded amphibolite derived from strongly deformed recrystallised pillow lava with flattened and extended dark pillow cores and pale pillow rims and inter-pillow quartz.

degree of alteration, flattened quartz in pillow matrices, and tectonically disrupted sheets and veins of recrystallised tonalite and quartz are set in a matrix dominated by carbonate (dolomite

and ankerite) and biotite (Fig. 11). These rocks were also formerly interpreted as pyroclastic felsic volcanic rocks and included in the ‘felsic forma-tion’ of Nutman (1986).

Fig. 8. Deformed and recrystallised banded iron formation. A layer of folded quartz boudins can be seen near the top of the photograph. Similar layers were previously interpreted as depositional conglomeratic structures (‘‘flat pebble conglomeratic structure’’of Nutman (1986), and Nutman et al. (1984).

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Fig. 9. Intensely deformed recrystallised chert and metadolerite dyke, tectonically transposed into the regional tectonic fabric.

4. Conclusions

This study, based on new field investigations and mapping, supports the general conclusions of Rosing et al. (1996) and strengthens them by presenting new geological maps of portions of the Isua greenstone belt. These maps reveal the loca-tion, structure and relationships of the main pro-toliths and their altered equivalents. The field investigations indicate that a major part of the actinolite – chlorite – talc – tremolite amphibolites and hornblende – garnet – biotite – carbonate am-phibolites were derived from basaltic pillow lava and related epiclastic rocks, because deformed volcanic structures, including pillows and pillow breccias, are widespread in all these rocks. All stages of deformation can be observed between well preserved pillow lava and pillow lava breccia structures and compositionally banded amphibo-lite schists. Likewise transitions can be followed, on a regional scale, between tonalitic gneiss with relict igneous textures to quartzo – feldspathic schists and mylonites that were previously inter-preted to be of felsic volcanic origin. The previ-ously mapped stratigraphy Nutman (1986) is of little value in understanding the geology of the Isua greenstone belt and is best abandoned. This

Fig. 10. Strongly deformed, recrystallised metasomatically altered pillow lava (black garnet – biotite – carbonate amphibolite, with lenses of white inter-pillow quartz), and strongly deformed, disrupted and recrystallised veins of tonalite (more continuous grey layers and lenses). This rock unit was previously interpreted by Nutman (1986) as derived from pyroclastic felsic volcanic rocks.

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Fig. 11. Strongly deformed, recrystallised metasomatically altered pillow lava (black carbonate with remnants of white inter-pillow quartz and disrupted quartz veins). This rock unit was previously interpreted as being derived from a sedimentary conglomerate by Bridgwater and McGregor (1974) and Dimroth (1982), and as being derived from pyroclastic felsic volcanic rocks by Allaart (1976), and Nutman (1986).

oldest known greenstone belt consists mainly of tectonically juxtaposed slices of complexly de-formed and recrystallised, basalt, ultramafic rocks and chert – BIF.

Acknowledgements

This work is part of the Isua Multidisciplinary Research Project (IMRP) led by Peter Appel and Stephen Moorbath, and supported by the Danish National Science Research Council, the Commis-sion for Scientific Research in Greenland, the Greenland Bureau of Minerals and Petroleum, and the Geological Survey of Denmark and Greenland. Their support is gratefully acknowl-edged. The content of the paper was presented at the EUG meeting in Strasbourg on 31 March 1999.

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Where the BIF is least deformed the layers of

quartz and magnetite are generally 0.2 – 1.0 cm

thick and these layers reflect deformed primary

layering that developed on or below the ocean

floor. However even the least deformed layering is

substantially modified by deformation, and most

layering was folded and extended parallel to the

plunge of the regionally dominant folds and

asso-ciated lineations (Figs. 2, 3b and 8). In some cases

individual layers of quartz were disrupted into

isolated boudins or fracture-bounded tabular

seg-ments during deformation that preceded the

fold-ing (Fig. 8). These structures have previously been

interpreted as sedimentary (‘flat pebble

conglom-eratic structure’; Nutman, 1986, p. 17 – 18).

Primary conglomeratic structures occur in a

chert – BIF horizon in the northeastern part of the

greenstone belt, in the southern part of the central

tectonic domain (Fig. 2). These conglomerates

include both oligomict conglomerate with round

pebbles of quartz in a matrix of mainly quartz,

biotite and garnet (‘round pebble conglomeratic

structure’ of Nutman, 1986, p. 18), and polymict

conglomerate containing pebbles of quartz and

meta-basalt in a matrix of quartz, biotite and

garnet (Appel et al., 1998).

However, most chert – BIF is intensely

de-formed and in the northeast part of the

green-stone belt (northeast part of Fig. 2) primary

layering is extensively transposed into a new

tec-tonic layering (Fig. 9) and much of the chert

comprises mylonite or recrystallised mylonite. In

the southwest part of the greenstone belt (Fig. 3b)

many thin layers of banded quartz schist and

mylonite could have been derived from either

primary chert or from quartz veins.

3.4. Ultramafic rocks

Ultramafic rocks are of two kinds:

anthophyl-lite-rich rocks and layered serpentinites associated

with metagabbro and metapyroxenite. Each kind

of ultramafic rock is confined to different tectonic

slices of the greenstone belt.

In the southwestern part of Fig. 3b,

anthophyl-lite-rich ultramafic rocks form thin layers within

actinolite – chlorite – talc – tremolite

amphibolite

with deformed pillow lava structures, and thicker

fault-bounded layers. The apparently simple,

mas-sive field appearance of these rocks results from

coarse radiating clusters of anthophyllite that

form most of the rocks. However this overprints a

complex history of older tectonic layering and

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J.S.Myers/Precambrian Research105 (2001) 129 – 141 137

deformed mafic and ultramafic dykes. These

ultra-mafic rocks could have originated as either

ko-matiite flows or sub-volcanic intrusions.

In the northeast part of Fig. 3b, ultramafic

rocks comprise serpentinite spatially associated

with compositionally layered metagabbro and

metapyroxenite. These ultramafic rocks appear to

be derived from layered dunite – peridotite –

pyrox-enite – gabbro intrusions.

3.5. Quartzo

-

feldspathic schist

Prominent layers of quartzo – feldspathic schist

(Figs. 2 and 3) have previously been interpreted as

derived from felsic volcanic rocks (Allaart, 1976;

Nutman, 1986). Most of these rocks are intensely

deformed schists or mylonites, but the amount of

Fig. 6. Banded amphibolite derived from strongly deformed recrystallised pillow lava with flattened and extended dark pillow cores and pale pillow rims.

Fig. 5. Deformed and recrystallised high-Mg basalt pillow lava with inter-pillow quartz. The pale spots within the pillows are deformed occelli.

deformation in the thickest layer in Fig. 3b

de-creases towards the east and the schist passes

gradationally into tonalitic gneiss. To the

north-west in Fig. 3b, thin attenuated and boudinaged

layers of quartzo – feldspathic schist and mylonite

occur within garnet – biotite – carbonate

amphibo-lite with discoidal bodies of quartz (Fig. 10).

These rocks were previously interpreted as

pyro-clastic felsic volcanic rocks (Allaart, 1976;

Nut-man, 1986), but are here interpreted as deformed

and recrystallised pillow lavas (with discoidal

quartz representing pillow matrices) and

de-formed, tectonically disrupted, sheets of

recrys-tallised tonalite, derived from veins intruded into

the pillow lavas, as these features can be observed

in the least strained parts of this unit. In other

places where the pillow lavas suffered a greater

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Fig. 7. Banded amphibolite derived from strongly deformed recrystallised pillow lava with flattened and extended dark pillow cores and pale pillow rims and inter-pillow quartz.

degree of alteration, flattened quartz in pillow

matrices, and tectonically disrupted sheets and

veins of recrystallised tonalite and quartz are set

in a matrix dominated by carbonate (dolomite

and ankerite) and biotite (Fig. 11). These rocks

were also formerly interpreted as pyroclastic felsic

volcanic rocks and included in the ‘felsic

forma-tion’ of Nutman (1986).

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J.S.Myers/Precambrian Research105 (2001) 129 – 141 139

Fig. 9. Intensely deformed recrystallised chert and metadolerite dyke, tectonically transposed into the regional tectonic fabric.

4. Conclusions

This study, based on new field investigations

and mapping, supports the general conclusions of

Rosing et al. (1996) and strengthens them by

presenting new geological maps of portions of the

Isua greenstone belt. These maps reveal the

loca-tion, structure and relationships of the main

pro-toliths and their altered equivalents. The field

investigations indicate that a major part of the

actinolite – chlorite – talc – tremolite

amphibolites

and hornblende – garnet – biotite – carbonate

am-phibolites were derived from basaltic pillow lava

and related epiclastic rocks, because deformed

volcanic structures, including pillows and pillow

breccias, are widespread in all these rocks. All

stages of deformation can be observed between

well preserved pillow lava and pillow lava breccia

structures and compositionally banded

amphibo-lite schists. Likewise transitions can be followed,

on a regional scale, between tonalitic gneiss with

relict igneous textures to quartzo – feldspathic

schists and mylonites that were previously

inter-preted to be of felsic volcanic origin. The

previ-ously mapped stratigraphy Nutman (1986) is of

little value in understanding the geology of the

Isua greenstone belt and is best abandoned. This

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oldest known greenstone belt consists mainly of

tectonically juxtaposed slices of complexly

de-formed and recrystallised, basalt, ultramafic rocks

and chert – BIF.

Acknowledgements

This work is part of the Isua Multidisciplinary

Research Project (IMRP) led by Peter Appel and

Stephen Moorbath, and supported by the Danish

National Science Research Council, the

Commis-sion for Scientific Research in Greenland, the

Greenland Bureau of Minerals and Petroleum,

and the Geological Survey of Denmark and

Greenland. Their support is gratefully

acknowl-edged. The content of the paper was presented at

the EUG meeting in Strasbourg on 31 March

1999.

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