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 textures 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:
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.
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 ‘amphibolite 9 garnet 9 carbonate 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
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 – tremolite amphibolite 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
. 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.
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.
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.
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
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.
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.
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
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.
Fig. 9. Intensely deformed recrystallised chert and metadolerite dyke, tectonically transposed into the regional tectonic fabric.
4. Conclusions
