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Morphology as an Indictor of Biogenicity for 3.5–3.2 Ga Fossil Stromatolites from the Pilbara Craton, Western Australia

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Stromatolites were recognised in 3.5–3.4Ga rocks from the Pilbara Craton 30 years ago (Walter et al. 1980; Lowe 1980), but their biogenicity has been cast in doubt by recent studies of abiogenic features with gross similarity to biogenic forms (Grotzinger and Rothman 1996; Garcia-Ruiz et al. 2003) and new interpretations of geological settings that differ from original models (see below). This paper reviews the geological setting of Pilbara fossils and assesses stromatolite biogenicity, showing the critical value of stromatolite morphology when used in combination with well documented geological context.
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Morphology as an Indictor of Biogenicity
for 3.5–3.2 Ga Fossil Stromatolites from
the Pilbara Craton, Western Australia
Martin J. Van Kranendonk
1 Introduction
Stromatolites were recognised in 3.5–3.4 Ga rocks from the Pilbara Craton 30 years
ago (Walter et al. 1980; Lowe 1980), but their biogenicity has been cast in doubt
by recent studies of abiogenic features with gross similarity to biogenic forms
(Grotzinger and Rothman 1996; Garcia-Ruiz et al. 2003) and new interpretations
of geological settings that differ from original models (see below). This paper
reviews the geological setting of Pilbara fossils and assesses stromatolite biogeni-
city, showing the critical value of stromatolite morphology when used in combina-
tion with well documented geological context.
2 Geological Setting of Proposed Fossil Localities
Proposed stromatolite fossils occur within the 3.53–3.165 Ga East Pilbara Terrane
(EPT) of the Pilbara Craton, Western Australia (Fig. 1: Van Kranendonk 2007;
Van Kranendonk et al. 2007a,b). The EPT consists of four unconformity-bound
volcano-sedimentary groups (Pilbara Supergroup) and related ultramafic-mafic and
granitic intrusive rocks (Fig. 2), all of which are unconformably overlain by rocks
of the 3.07–2.94 Ga De Grey Supergroup, deformed and intruded by 3.07–2.83 Ga
granitic rocks, and then unconformably overlain by the 2.78–2.63 Ga Fortescue
Group (Van Kranendonk et al. 2007a). The lower three groups of the Pilbara
M.J. Van Kranendonk (*)
Geological Survey of Western Australia, 100 Plain St., East Perth, WA 6004, Australia
Geographical Sciences, The University of Western Australia, 35 Stirling Hwy, Crawley,
WA 6009, Australia
School of Earth and Environment, The University of Western Australia, 35 Stirling Hwy,
Crawley, WA 6009, Australia
e-mail: vankranendonk@dmp.wa.gov.au
J. Reitner et al., Advances in Stromatolite Geobiology,
Lecture Notes in Earth Sciences 131, DOI 10.1007/978-3-642-10415-2_32,
#Springer-Verlag Berlin Heidelberg 2011
537
Supergroup are dominantly volcanic, and were deposited from 3.53 to 3.24 Ga as
the result of discrete mantle plume events, whereas the c. 3.2–3.165 Ga Soanesville
Group is largely composed of clastic sedimentary rocks and is interpreted to
relate to rifting of EPT margins, prior to the onset of terrane accretion at 3.07 Ga
(Van Kranendonk et al. 2007a,2010).
Strong claims for early life have been described from several stratigraphic
horizons within the Pilbara Supergroup, including (Fig. 2):
1. c. 3.48 Ga Dresser Formation (diverse stromatolites, putative microfossils,
highly negative d
13
C on carbonaceous material, sulfur isotopes) (Walter et al.
1980; Ueno et al. 2004; Van Kranendonk 2006; Philippot et al. 2007)
Fig. 1 Geological map of the central East Pilbara Terrane, showing main lithological units and
localities mentioned in the text
538 M.J. Van Kranendonk
2. c. 3.46 Ga Apex chert (putative microfossils, Laser-Raman spectrography)
(Schopf 1993; Schopf et al. 2007)
3. c. 3.40 Ga Strelley Pool Chert (diverse stromatolites, Carbon isotope data, and
organic geochemistry) (Lowe 1980; Hofmann et al. 1999; Allwood et al. 2006;
Van Kranendonk 2006; Marshall et al. 2007)
4. c. 3.35 Ga Euro Basalt (ichnofossils in altered basaltic glass) (Banerjee et al.
2007; Furnes et al. 2007)
5. c. 3.24 Ga Sulphur Springs Group (putative microfossils in sulphides) (Rasmussen
2000)
6. c. 3.20 Ga Soanesville Group (organic matter, highly negative d
13
C on carbona-
ceous material, putative microfossils) (Duck et al. 2007; Glikson et al. 2008)
Fig. 2 Stratigraphic column of the Pilbara Supergroup, showing early life localities
Morphology as an Indictor of Biogenicity for 3.5–3.2 Ga Fossil Stromatolites 539
Previous models of the geological setting of the units mentioned above include
(1) quiet, shallow water, evaporative lagoon (Groves et al. 1981; Walter et al.
1980); (2) marine clastic sandstone (Schopf 1993); (3) sabkha to shallow marine
environment (Lowe 1980,1983); (4) submarine pillow basalt (Banerjee et al. 2007;
Furnes et al. 2007); (5) active chemoautotrophic microbial community within a
hydrothermal vent (Rasmussen 2000); (6) sediment-hosted microbial community
associated with active hydrothermal activity (Duck et al. 2007).
Current work has indicated that the geological settings for the Dresser Formation
and Apex Chert are significantly different from those previously proposed, whereas
detailed mapping of the Strelley Pool Chert has largely confirmed previous inter-
pretations, although across a much broader area.
Detailed geological mapping, in combination with petrology and studies of
alteration mineralogy and geochemistry, have shown that the Dresser Formation
was deposited within an active volcanic caldera affected by syn-volcanic growth
faulting, extensive hydrothermal fluid circulation, and eruption of minor felsic tuff
(Nijman et al. 1998; Van Kranendonk 2006; Van Kranendonk and Pirajno 2004;
Van Kranendonk et al. 2008). In detail, three main periods of deposition have been
recognised (1) an early shallowing up succession of minor basal conglomerate,
and laminated siderite, culminating with rippled carbonate sediments that contain
widespread stromatolites and at least local evidence for subaerial exposure; (2) a
fining upward succession of conglomerate grading up to laminated cherts (silicified
carbonates) that contains local stromatolites; (3) a fining up succession of conglom-
erate and sandstone passing up to laminated cherts (silicified carbonates) (Van
Kranendonk et al. 2008). Each of these three successions is marked by a basal
unconformity. Whereas the lower two successions contain thick sills and veins of
coarsely crystalline barite and chert, the upper succession lacks barite and contains
thick sandstone deposits derived from volcanic sources, including a 10 cm thick
unit of felsic volcaniclastic sandstone dated at 3,481 2 Ma (Van Kranendonk
et al. 2008). Growth faulting allowed penetration of seawater deep into the crust,
where it was heated and expelled back up to the surface, causing steam-heated acid-
sulphate alteration of the near-surface footwall basalts (Van Kranendonk and Pirajno
2004). Leaching of Ba, Si and Fe from footwall basalts combined with magmatic
S and seawater sulphate to precipitate barite, silica, and sulphides (pyrite and
sphalerite) within fossil hydrothermal veins (Fig. 3). The sedimentary units were
affected by growth faulting prior to eruption of disconformably overlying high-Mg
pillow basalts.
Samples with putative microfossils from the c. 3.46 Ga Apex Chert were
previously interpreted to be marine sandstones (Schopf 1993). However, recent
mapping and geochemical data revealed that the samples were collected from
a fossil hydrothermal breccia vein in the subsurface to bedded sandstone
(Brasier et al. 2002,2005; Van Kranendonk 2006). This, in combination with
more highly detailed petrography and geochemistry on the samples, has cast
considerable doubt on the biogenicity of the putative microfossils (Brasier et al.
2002,2005), although other studies continue to support a biogenic origin of these
structures (Schopf et al. 2007), and it has been recognised that microfossils
540 M.J. Van Kranendonk
Fig. 3 Dresser Formation model of deposition within an active volcanic caldera affected
by hydrothermal fluid circulation. Downward circulating seawater gets heated up and interacts
with reducing fluids and volatiles from a degassing magma chamber. These fluids leach Ba, Si and
Fe from footwall basalt, mix with seawater containing dissolved sulphate and are then expelled
to the surface along syn-volcanic growth faults where they precipitate silica, barite and pyrite
(from Van Kranendonk 2006)
Morphology as an Indictor of Biogenicity for 3.5–3.2 Ga Fossil Stromatolites 541
could occur within a subsurface hydrothermal vein through three different ways
(Van Kranendonk 2007).
The Strelley Pool Chert was deposited across a regional angular unconformity
on older rocks during a marine transgression from subaerial exposure to fully
marine conditions (Allwood et al. 2006; Van Kranendonk 2006,2007). Members
of the formation include: basal fluviatile to beach and nearshore clastic deposits
(conglomerate and well-sorted quartz sandstone), up to 1 km thick in places; a
middle member of shallow marine, grading up to supratidal, carbonates, typically
8 m thick; a disconformably overlying member of coarse clastic, submarine alluvial
fans up to 15 m thick (Fig. 4). This sedimentary formation is conformably overlain
by pillow basalt with local, thin, units of felsic tuff, one of which has returned a
U-Pb SHRIMP age of 3,350 2 Ma (Van Kranendonk et al. 2007a).
3 Stromatolite Morphology as an Indicator of Biogenicity
Stromatolites occur at several stratigraphic levels within the Dresser Formation and
have diverse forms, even within continuous horizons. Significantly, although most
surface outcrops are of chert, rare outcrop and drillcore samples show that the
Fig. 4 Outcrop sketch of part of the Strelley Pool Formation, where it lies unconformably on
c. 3.51 Ga, folded rocks of the Coonterunah Subgroup (Warrawoona Group). Note three main
members relating to depositional environment
542 M.J. Van Kranendonk
primary material was marine carbonate (ankerite and siderite), replaced by later
silica. Stromatolite forms include: coniform, columnar, domical and wrinkly mats
in shallow water, rippled marine carbonate sediments (Fig. 5), suggesting possible
Fig. 5 Stromatolites in bedded carbonate sedimentary rocks of the c. 3.48 Ga Dresser Formation.
(a) Wrinkly laminated stromatolite mat overlying rippled carbonate sandstone (arrow). (b)
Wrinkly laminated stromatolite mat overlain by domical stromatolites. (c) Oblique bedding
plane view of coniform stromatolites, which lie within rippled carbonate sandstone. (d) Oblique
bedding plane view of irregular-shaped stromatolites in silicified, but finely laminated and
rippled, carbonate sedimentary rock. (e) Oblique bedding plane view of 30-cm diameter domical
stromatolite with wrinkly laminated flanks and smooth core. (f) Cross-sectional outcrop view of
small, domical stromatolite within shallow-dipping carbonate sedimentary rocks that have been
partly replaced by barite and silica: note the high amplitude of the column and the flat underlying
and overlying strata
Morphology as an Indictor of Biogenicity for 3.5–3.2 Ga Fossil Stromatolites 543
photoautotrophy; domical and wrinkly laminated forms at, and downflow from,
hydrothermal vents (Fig. 6), suggesting chemoautotrophy; and wrinkly laminates
within some hydrothermal veins, parallel to vein walls, suggesting subsurface
chemoautotrophs. Key morphological indicators of stromatolite biogenicity are
coniform shapes in rippled carbonate sand (Fig. 5c), wrinkly lamination within
coniform stromatolites (Fig. 7a) and laminated mats (Fig. 7b, c), microbially bound
ripples in carbonate sand (Fig. 7d), irregular domical forms within 20 cm along
strike of coniform varieties (Figs. 5c, d), and the presence of draped sedimentary
wedges off domical stromatolite flanks, indicative of stromatolite growth during
sedimentation (Van Kranendonk 2006). The occurrence of stromatolites within
particulate carbonate sediment indicates the laminations are not carbonate crusts.
Stromatolites are truncated by flat pebble conglomerates, indicating formation
during sedimentation and not as a result of post-depositional diagenetic effects.
The variation along strike from coniform to irregular shaped stromatolites within
rippled sands with flat pebble conglomerates, suggests interplay between growth
morphology of stromatolites and water currents in very shallow, to supratidal,
conditions, in much the same way as in modern day Shark Bay (e.g. Groves et al.
1981). Thin sections show that the morphology of small (<1 cm) columnar
stromatolites in pyrite-replaced carbonate sediment is independent of the diagenetic
growth of barite crystals (Fig. 8), also countering a post-depositional mode of
formation. These structures are characterised by very fine-scale, wavy laminations
within relatively high amplitude columns (A in Fig. 8), and by laminae that either
terminate at column margins (no growth wall: left of B in Fig. 8), or are bound
within a smooth growth wall (C in Fig. 8). Locally, individual columns are joined
by bridging laminae (D in Fig. 8). Stromatolite biogenicity is supported by sulphur
Fig. 6 Stromatolites at, and adjacent to, syn-depositional hydrothermal veins in the c. 3.48 Ga
Dresser Formation, Pilbara Craton. (a) Low-amplitude domical stromatolite (s) at the mouth of a
hydrothermal vein (v), within sedimentary rocks that have been extensively replaced by barite.
Note the massive and smooth-weathering nature of vein barite (v) where it cuts bedding (b), vs. the
more wrinkly laminated character of the barite-replaced sedimentary rocks. (b) Large, low-
amplitude domical stromatolite, approximately 5 m along strike of (a), showing flat bedding
overlying stromatolite and extensive replacement of primary sediment by hydrothermal barite
(light grey areas)
544 M.J. Van Kranendonk
Fig. 7 Key morphological indicators of biogenicity in Dresser Formation stromatolites. (a) Cross-
sectional outcrop view of fine-scale, wrinkly lamination within coniform stromatolite. (b) Cross-
sectional outcrop view of wrinkly laminated stromatolite mat, directly overlying rippled carbonate
sedimentary rock (width of view ~30 cm). (c) Oblique bedding plane view of wrinkly laminated
stromatolite mat, directly overlying rippled carbonate sedimentary rock, showing numerous
circular forms (lenscap is ~3 cm). (d) Cross-sectional outcrop view of wrinkly lamination within
coniform stromatolite that is directly overlain by microbial bound rippled carbonate sedimentary
rock with a small domical stromatolites (arrow)
Fig. 8 Plane light thin section view of columnar, pyrite-replaced carbonate stromatolites, from
95.35 m in diamond drillhole PDP2c through the Dresser Formation (width of view ~3 cm: Van
Kranendonk et al. 2008). These structures show high amplitude columnar forms in flat-bedded
sedimentary rocks that have wrinkly internal laminations (A) and either ragged margins (B) or
smooth growth walls (C), with bridging laminae between individual columns (D)
Morphology as an Indictor of Biogenicity for 3.5–3.2 Ga Fossil Stromatolites 545
Fig. 9 Stromatolites in the marine carbonate member of the Strelley Pool Chert. (a) Oblique
bedding plane view of relatively uniform size coniform stromatolites in silicified dolomite. (b)
Cross-sectional outcrop view of coniform stromatolites in dolomite, showing near-isopachous
laminations but varying form, including incipiently branching form on the left: note the start of
coniform stromatolite from horizontally laminated dolomite at arrow. (c) Cross-sectional outcrop
view of coniform stromatolite in partly silica-replaced dolomite, with two branching forms on left-
hand side (just to right of pen). (d) Cross-sectional outcrop view of small, branched columnar
stromatolite in partly silica-replaced dolomite (width of view ~5 cm). (e) Cross-sectional outcrop
view of microbially-bound ripples in partly silica-replaced carbonate sandstone, which occurs
along strike of coniform stromatolites (lenscap ~3 cm). (f) Cross-sectional outcrop view of
coniform stromatolite, showing non-isopachous laminae and flat overlying bedding (width of
cone base ~4 cm)
546 M.J. Van Kranendonk
isotope data, carbon isotope data from clots of carbonaceous material in micritic
carbonate, by possible microbial remains and highly negative d
13
C on kerogen and
methane in fluid inclusions in hydrothermal veins (Ueno et al. 2001,2004,2006;
Glikson et al. 2008; Van Zuilen et al. 2008).
Stromatolites in the Strelley Pool Chert have diverse and unique morphologies in
different depositional settings: shallow marine dolomites contain conical, incipi-
ently branching conical, and branched conical forms, columnar and branching
columnar forms, and wavy laminate mats binding rippled carbonate sand (Fig. 9);
supratidal carbonates with low-amplitude, locally windblown, ripples contain low
amplitude domes and wrinkly laminated mats (Fig. 10); clastic rocks contain
Fig. 10 Stromatolites in the
supratidal carbonate member
of the Strelley Pool Chert. (a)
Oblique bedding plane view
of wind-blown ripples in
finely-layered silicified
dolomite. (b) Cross-sectional
outcrop view of small
domical stromatolites in
silicified dolomite: note
horizontal overlying and
underlying beds (Width of
view ~12 cm). (c) Oblique
bedding plane view of top of a
small domical stromatolite in
silicified, finely-laminated
dolomite
Morphology as an Indictor of Biogenicity for 3.5–3.2 Ga Fossil Stromatolites 547
laminar mats with abundant clotty kerogen and clasts of laminated carbonaceous
material (Fig. 11: Van Kranendonk 2007).
Key morphological indicators of biogenicity include a diverse array of macro-
scopic and microscopic features (Hofmann et al. 1999; Van Kranendonk et al. 2003;
Fig. 11 Stromatolites in the clastic member at the top of the Strelley Pool Chert. (a) Cross-
sectional outcrop view of sandstone with planar, black stromatolite mats: note trapped sand grains
(light dots) in lower two mats and in mat just above pen, and also the cross-cutting chert vein. (b)
Plane polarized light thin section view of planar, black stromatolite mat in sandstone from (a),
showing the presence of clotted carbonaceous material and trapped sand grains (width of view is
2.5 cm). (c) Plane polarized light thin section view of clast of laminated carbonaceous material in
sandstone with volcanic clasts (width of view ~5 mm). (d) Plane polarized light thin section view
of clotty carbonaceous material, with sand grains and silica-filled gas bubbles (Width of view
~2 mm)
548 M.J. Van Kranendonk
Allwood et al. 2006; Van Kranendonk 2006,2007). In marine dolomites, Van
Kranendonk et al. (2003) have convincingly demonstrated that the laminae that
define most of the macroscopic stromatolites, although locally isopachous (Fig. 9b),
in reality show a variety of features and are not typical of seafloor crusts (Fig. 9f).
Specifically, evidence for microbial binding of carbonate sediment has been
observed (Fig. 9e) and it has also been observed that coniform stromatolites
occur along strike of rippled sediment, indicating stromatolite growth in moving
water with loose, particulate carbonate sand (Fig. 12: Van Kranendonk et al. 2003).
Evidence for axial growth zones within some large coniform stromatolites has also
been reported (Grey 2008).
The most compelling morphological features of biogenicity includes a high-
amplitude coniform stromatolite that grew upwards through periods of alternating
low energy and high energy depositional events (Fig. 13a). Another convincing
piece of morphological evidence is from domical stromatolites in supratidal carbo-
nates that show truncated laminae along a smooth growth wall (1 on Fig. 13b),
onlapping laminae from one column to the next (2 on Fig. 13b), infill of the
intercolumn space by particulate carbonate (3 on Fig. 13b), and truncation of
stromatolite growth by a period of erosion and subsequent clastic sediment deposi-
tion (4 and sst on Fig. 13b). A final piece of evidence is the presence of different
stromatolite morphologies within close proximity along strike (Fig. 13c) and
different size conical forms on individual bedding planes (Fig. 13d, e).
Fig. 12 Evidence for stromatolite growth during sediment accumulation in the marine carbonate
member of the Strelley Pool Chert: (a) Cross-sectional outcrop view of silicified coniform
stromatolite flank, with onlapping sediment layers (arrows;bars are in centimetres); (b) Oblique
cross-sectional outcrop view of coniform stromatolite (c) in silicified carbonate sandstone with
evidence of deposition under high-energy conditions, including ripples (circle), small channel fill
of flat pebble conglomerate (above short dashed line), and high-amplitude sedimentary swale (top
2/3 of image, between upper and lower heavy dashed lines) with internal unconformities (middle
heavy dashed line)
Morphology as an Indictor of Biogenicity for 3.5–3.2 Ga Fossil Stromatolites 549
Fig. 13 Key morphological indicators of biogenicity in the Strelley Pool Chert. (a) Cross-
sectional outcrop view of high-amplitude coniform stromatolite in dolomite, passing up from
quiet water laminates, through a unit of flat-pebble conglomerate deposited during a high-energy
(storm?) event, and back into quiet water laminates. (b) Plane polarized light thin section view of
small columnar stromatolites from the supratidal carbonate member (Fig. 10b), showing smooth
growth wall that truncates internal laminations in the right-side stromatolite (1), laminations of the
left-side stromatolite that lap onto the growth wall of the adjacent column (2), sedimentary infill of
the intercolumn space (3), and capping transgressive sand bed (sst) that truncates and erodes
550 M.J. Van Kranendonk
Several additional features support the contention that stromatolite morphology
is not the result of post-depositional diagenesis and/or growth of crystal splays,
including the fact that crystal splays locally nucleate on the top of coniform
stromatolites (Fig. 13f). Locally, it has been observed that horizons of crystal splays
terminate the growth of stromatolites and are overlain by thinly bedded, rippled,
supratidal carbonate sediments, which combines to suggest that the crystal splay
beds formed during periods of subaerial exposure when stromatolite growth ceased
(Fig. 14). Additional outcrop-scale evidence for biogenicity comes from the distri-
bution of stromatolites along and across strike (Van Kranendonk 2007; Allwood
et al. 2006). The best example of this is at the Trendall locality, where stromatolite
size and shape varies along strike in three onlapping biostromes, from larger and
more complex forms with greater inheritance (branching and incipiently branching
forms) near the upslope part of the bioherm, to smaller and more simple forms
downslope at the toe of the bioherms (Fig.14).
Stromatolite biogenicity is supported by highly negative carbon isotope data and
organic geochemistry on kerogenous material in clastic sedimentary rocks from
these outcrops (Marshall et al. 2007).
4 Summary
Convergent lines of evidence from investigations of putative fossil assemblages
from several distinct horizons in the East Pilbara Terrane of the Pilbara Craton point
to abundant and diverse microbial life on early Earth. A key indicator of biogenicity
in all studied examples is morphology, but this must be set within a well constrained
geological context.
ä
Fig. 13 (Continued) laminae in the underlying stromatolites (4) (height of view ~2 cm). (c) Cross-
sectional outcrop view of laminated dolomite showing large, incipiently branching stromatolite
and small, branched columnar stromatolite on flank at top right (detail shown in Fig. 9d). (d)
Highly oblique bedding plane view of large (8 cm), slightly asymmetrical columnar stromatolite in
partly silicified dolomite, flanked by adjacent, smaller coniform stromatolites. (e) Bedding plane
view of elongate coniform stromatolites with varying size. (f) Cross-sectional outcrop view of
small coniform stromatolite capped by a radiating crystal splay
Morphology as an Indictor of Biogenicity for 3.5–3.2 Ga Fossil Stromatolites 551
Fig. 14 Outcrop map of part of the Strelley Pool Chert, preserved in cross-section, showing carbonate-encrusted boulders of the basal clastic member,
onlapping stromatolitic biostromes (purple) and diagenetic crystal splays within carbonate member, domical stromatolite horizons within the silicified
supratidal carbonate member (red lines in light blue unit), and overlying clastic member comprising five fining-upward units interpreted as the depositional
products of receding submarine fans: red star denotes occurrence of laminar stromatolites within upper clastic member, shown in Fig. 11 (from Van
Kranendonk 2007)
552 M.J. Van Kranendonk
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... The transition between fluvial to tidedominated shelf settings are reported from some of the Mesoarchean siliciclastic successions (Eriksson et al., 1994(Eriksson et al., , 2005Eriksson and Simpson, 2000). Well preserved alluvial fan to fan-delta deposits are reported from the >3.44 Ga Hooggenoeg Formation of the Fig Tree Group (Nocita and Lowe, 1990;Eriksson et al., 2006;Heubeck et al., 2013;Drabon et al., 2017;Heubeck, 2019) in the Barberton Greenstone Belt, Kaapvaal craton and from the lower parts of the ~3.55 Ga siliciclastics of the Warrawoona Group, Pilbara craton (Buick et al., 1995;Van Kranendonk, 2011). Fan-delta to shelf sedimentation record became increasingly apparent from the Mesoarchean successions, e.g. ...
... The transition between fluvial to tidedominated shelf settings are reported from some of the Mesoarchean siliciclastic successions (Eriksson et al., 1994(Eriksson et al., , 2005Eriksson and Simpson, 2000). Well preserved alluvial fan to fan-delta deposits are reported from the >3.44 Ga Hooggenoeg Formation of the Fig Tree Group (Nocita and Lowe, 1990;Eriksson et al., 2006;Heubeck et al., 2013;Drabon et al., 2017;Heubeck, 2019) in the Barberton Greenstone Belt, Kaapvaal craton and from the lower parts of the ~3.55 Ga siliciclastics of the Warrawoona Group, Pilbara craton (Buick et al., 1995;Van Kranendonk, 2011). Fan-delta to shelf sedimentation record became increasingly apparent from the Mesoarchean successions, e.g. ...
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This paper describes the results of detailed mapping and combined structural and sedimentologic investigation of the similar to 3.5 Ga old North Pole Chert along the eastern and southern flanks of the North Pole granitoid dome of the Pilbara craton. The following relationships have been found. (1) Stepwise and repeated thickening over north block-down listric normal growth faults, now tilted to 120 degrees E. Apart from the thickness distribution, the growth fault character has been inferred from progressive upward decrease of throw towards gentle flexuring, and the presence of roll-over anticlines, in a restored pre-tilt cross-section. The faults are unrelated to the actual geometry of the dome, and overprinted by pre-doming compressive deformation. (2) A reconsideration of the sedimentary model of the chert-barite unit with emphasis on vertical grainsize distribution and facies architecture shows direct synsedimentary fault control. (3) Black chert veins, clustered in the hangingwall blocks of the growth faults follow a conjugate fracture pattern in and immediately beneath the North Pole Chert, symmetrical with respect to the synsedimentary fault geometry. Upwards, the veins terminate in the lowermost chert-barite unit of the North Pole Chert; downwards, the majority of the veins converge towards centres at the growth faults in the underlying basalt. (4) Barite occurs: in veins, in association with black chert; as large synsedimentary mounds (15 m high, 50 m wide), formed on the original shallow subaqueous basin floor; as commonly silicified sinters and sedimentary or early diagenetic evaporite associated with littoral and/or stromatoloidal facies. The barite mounds are directly connected with chert-barite veins, and are uplapped by banded sedimentary chert and diamictite with barite clasts. Field relationships suggest precipitation of barite, particularly in the mounds, may have been primary and synsedimentary. Vein barite at or near the original depositional surface of the chert is associated with base metal sulphide at depth. The tectono-sedimentary relationships between normal faults, sedimentation patterns and thick:ness distribution, primary and secondary chert, synsedimentary barite mounds, stratiform barite and chert-barite veins, support a tensional fault-controlled, hydrothermal model with Ba, Si and sulphide emanation in ca 50 m depth of water. Such a model envisages boiling and/or degassing above vents to locally mix a normally stratified water body, causing instantaneous oxidization into sulphate. The sedimentary model of the North Pole Chert indicates repeated uplift and subsidence, with development of faults and fractures, the intensity of which diminished from the lowermost chert-barite unit of the North Pole Chert to the uppermost, fifth unit. (C) 1999 Elsevier Science B.V.
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More than 600 specimens of ∼3.5 Ga-old hydrothermal silica dikes from the North Pole area, Pilbara craton, Western Australia, have been studied petrographically. The kerogens in 44 samples have been analyzed isotopically (C and N) and chemically (C, N, and H). The silica dikes are composed mainly of fine-grained silica (modal abundance: >97%) and are classified into two types by minor mineral assemblages: B(black)-type and G(gray)-type. The B-type silica dikes contain kerogen (0.37 to 6.72 mgC/g; average 2.44 mgC/g, n = 21) and disseminated sulfides, dominantly pyrite and Fe-poor sphalerite. In some cases, carbonate and apatite are also present. Their silica-dominated and sulfide-poor mineral assemblages suggest precipitation from low-temperature reducing hydrothermal fluid (likely 100–200°C). On the other hand, the G-type silica dikes are sulfide-free and concentrations of kerogen are relatively low (0.05 to 0.41 mgC/g; average 0.17 mgC/g, n = 13). They typically contain Fe-oxide (mainly hematite) which commonly replaces cubic pyrite and rhombic carbonate. Some G-types occur along secondary quartz veins. These textures indicate that the G-type silica dikes were formed by postdepositional metasomatism (oxidation) of the B-types, and that the B-types probably possess premetasomatic signatures. The δ13C values of kerogen in the B-types are −38.1 to −33.1‰ (average −35.9‰, n = 21), which are ∼4‰ lower than those of the G-types (−34.5 to −30.0‰; average −32.2‰, n = 19), and ∼6‰ lower than bedded chert (−31.2 to −29.4‰; average −30.5‰, n = 4). This indicates the preferential loss of 12C during the metasomatism (estimated fractionation factor: 0.9985). Considering the metasomatic effect on carbon isotopes with probably minor diagenetic and metamorphic overprints, we conclude that the original δ13C values of the kerogen in the silica dikes would have been heterogeneous (∼5‰) and at least some material had initial δ13C values of ≤ −38‰. The inferred 13C-depletions of organic carbon could have been produced by anaerobic chemoautotrophs such as methanogen, but not by aerobic photoautotrophs. This is consistent with the estimated physical and chemical condition of the hydrothermal fluid, which was probably habitable for anaerobic and thermophilic/hyperthermophilic chemoautotrophs. Alternatively, the organic matter may have been possibly produced by abiological reaction such as Fischer-Tropsch Type (FTT) synthesis under the hydrothermal condition. However, the estimated condition is inconsistent with the presence of the effective catalysts for the FTT reaction (i.e., Fe-Ni alloy, magnetite, and hematite). These lines of evidence suggest the possible existence of biosphere in the ∼3.5 Ga sub-seafloor hydrothermal system.
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The 3.4 Ga-old Strelley Pool Chert is a 25-m thick sedimentary unit near the top of the predominantly volcanic Warrawoona Group in greenstone belts of the eastern Pilbara Block, Western Australia. It is here subdivided into 5 members containing 13 lithofacies. The basal Member, I, is composed of quartzose sandstone deposited in a high-energy wave- or tide-dominated shallow-water system. Overlying this are Members II and III, which make up the bulk of the formation and were deposited in a low-energy, partially restricted hypersaline basin. They record a predominantly regressive succesion of deposits including subaqueous laminite, stromatolite and evaporite; stromatolite, carbonaceous laminite, black-and-white banded chert, evaporite and intraformational detrital units deposited under intermittently to predominantly exposed conditions; and subaerially deposited windblown sand, evaporite and evaporite-solution layers. Members IV and V record the progradation of a volcaniclastic alluvial fringe.
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A sequence of c. 3.2Ga low-grade sedimentary rocks, bimodal volcanic rocks, and subvolcanic layered mafic-ultramafic sills and dykes occur in greenstone belts that flank the margins of the 3.51–3.24Ga East Pilbara Terrane of the Pilbara Craton, Western Australia. U–Pb SHRIMP zircon data and Sm–Nd model age data suggest that these rocks represent juvenile addition of material to the margins of the craton during a period of major extension, accompanied by the local emplacement of granitic rocks. This extensional event is also represented in adjacent terranes, by thick basalt flows with c. 3.2Ga zircon and Sm–Nd model ages, and by granitic rocks.Based on the available age data, geochemistry of basaltic rocks, and the distribution of this event in space and time within the craton, we interpret the c. 3.2Ga rocks to reflect widespread rifting of the protocraton and associated thinning of the lithosphere. This event resulted in separation of the flanking Kurrana and Karratha Terranes from their parent East Pilbara Terrane, which represents the ancient nucleus of the craton. These terranes were then reassembled during subsequent accretion at 3.07–2.905Ga. Combined, the data suggest the preservation of a Mesoarchean Wilson cycle in the Pilbara Craton.