Surface outcrops of the Tumbiana Formation display abundant and diverse stromatolites, ranging in size from a metre to centimetre scale (Fig. 2a). These overlie a thick section (up to hundreds of metres) of volcaniclastic sandstone and accretionary lapilli beds (Mingah Member), which in turn overlie thick subaerial basalt flows of the Kylena Formation. The Tumbiana Formation is overlain by thick, commonly vesicular subaerial basalt flows of the Maddina Formation. It was deposited at 2724 ± 5 Ma, as indicated by zircon U-Pb dating [3], either in a shallow marine or lacustrine environment associated with influx of riverine fresh water derived from the continents [4], [9], [42] and [52]. A maximum temperature of rock equilibration of about 300 °C has been estimated using metamorphic mineral assemblages and the degree of organization of organic matter on the PDP1 drill core samples [28].

In stromatolites, a few putative filamentous microfossils [45] and their palimpsests were interpreted as traces of phototrophic microbes [45] and [65]. In addition, 2α-methylhopane molecules that can be attributed to cyanobacteria [6] or an anoxygenic phototroph [40] were also detected in the Tumbiana Formation. It was thus proposed that these stromatolites formed primarily by the activity of photosynthesizers [65]. However, recent investigations of the carbon stable isotopes from different generations of carbonaceous material present in other Pilbara rocks analyzed for their molecular fossils were interpreted to indicate that the soluble hydrocarbons entered the rocks some 500 million years after deposition and therefore are not indigeneous to the original host rock [41].

Sedimentary organic matter strongly depleted in the carbon isotope ¹³C down to −60‰ has long been attributed to growth of aerobic, methanotrophic bacteria [20]. However, Hinrichs [22] showed that anaerobic methanotrophy involving sulfate as the electron acceptor according to the reaction CH₄ + SO₄ ²⁻ = HCO₃ ⁻ + HS⁻ + H₂O could be a viable mechanism for explaining the observed depletion of ¹³C in sedimentary organic carbon. According to this model, the 2.7 Ga organic carbon isotope excursion could indicate that sulfate concentrations were high enough for utilization by sulfate-reducing bacteria. In either case, the isotopic signature provides strong evidence for the existence of archaeal methanogens.

The Dresser Formation consists of interbedded units of chert–carbonate–barite and basalt that form a ring of hills which dip shallowly away from the North Pole Monzogranite in the core of the North Pole Dome. The monzogranite was emplaced into the core of the dome as a subvolcanic laccolith during eruption of the overlying, c. 3.46–3.42 Ga Panorama Formation [58]. The lowermost chert unit is intercalated with several barite (BaSO₄) lenses and is overlain by silicified felsic volcanogenic, hydrothermal breccia and bedded carbonate sediments that were deposited in a shallow water environment [27], [35] and [64]. This succession of chert, barite, volcanoclastic sandstone, hydrothermal breccia and carbonate is herein referred to as the chert–barite unit. This unit is overlain by pillowed basalt and underlain by spinifex-textured metabasalt that experienced low-grade metamorphism of between 100 and 350 °C [48]. The underlying komatiitic basalt is pervasively affected by intense hydrothermal alteration and transected by a myriad of veins–varying from centimeters to kilometers long–of barite and black and white cherts [21], [54], [56] and [60].

Pb–Pb isochron age of 3490 Ma was obtained from galena in barite of the chert–barite unit at North Pole [51]. This is in good agreement with a zircon U-Pb age of 3481 Ma recently obtained from a volcanoclastic layer of the PDP2 drill core [64] and a Sm–Nd age of 3480 Ma determined from volcanic, sedimentary and hydrothermal rocks of the chert–barite deposit and underlying volcanics [49].

Three models have been proposed for the setting of the chert–barite unit. One model suggests that the rocks were deposited as bedded carbonate–gypsum evaporites and mafic volcanogenic sediments in a quiet, shallow water marine lagoon separated from the open ocean by a sand bar [8] and [19]. In this model, the chert units were interpreted as silicified carbonates and bedded barite units as evaporative gypsum that were replaced by barite during early hydrothermal fluid circulation as attested by the extensive network of silica ± barite veins that underlie the formation. Subsequent hydrothermal circulation has been related either to structural doming [21], or rifting at 2.7 Ga during deposition of the Fortescue Group [7]. A second model suggests that the chert–barite units were deposited on oceanic crust during off-axis hydrothermal circulation and that the regional stratigraphy represents a set of thrust-bound ophiolite slices [24] and [45]. A third model suggests that the protoliths of the chert–barite unit were deposited during hydrothermal circulation in a volcanic caldera [35], [58], [59] and [60]. In this model, the silica and barite veins that underlie the bedded chert–carbonate units are interpreted as fossil hydrothermal fluid pathways that formed synchronously with, and were the cause of, deposition of the bedded chert–barite unit. Support for this interpretation arises from recent ¹⁴⁷Sm–¹⁴³Nd data indicating isochronous relationship between volcanic, sedimentary and hydrothermal rock types, including barite, at about 3.481 Ga [49]. In addition, bulk and individual fluid inclusion analysis led to the recognition of three main types of fluid populations preserved in intra-pillow quartz pods [15] and underlying hydrothermally-altered komatiitic basalts [39]: a metal depleted fluid, a Ba-rich and S-depleted fluid, and a Fe–S-rich end-member. The Cl/Br ratio of the metal depleted fluid inclusions (630) is similar to that of modern seawater (649) and has been interpreted as a remnant of Archean “seawater” [15]. In contrast, Ba- and Fe-rich brines have Cl/Br ratios (350 and 390) close to that of the bulk Earth value (420), hence arguing for a hydrothermal origin of these fluids.

Silica veins underlying the chert–barite unit contain small amounts of sulfides, carbonaceous filaments [14] and CH₄-bearing fluid inclusions with δ¹³C values between −30 and −35‰ [58] and as low as −56‰ [57], respectively. The carbonaceous filaments have been interpreted as microfossils of chemoautotrophs [58], whereas the associated ¹³C-depleted inclusion fluids are thought to have been produced by methanogens in surficial environments and recycled in the silica veins by hydrothermal fluids [57]. It is from similar intrusive silica veins beneath the 3460 Ma Apex chert that Schopf et al., [44] and [46] interpreted in situ carbonaceous material as “cellular microfossils”, whereas Brasier et al. [5] and Lindsay et al. [30] advocated a non-biogenic origin associated with Fischer–Tropsch synthesis for this material. The chert–barite unit contains the oldest known occurrence of stromatolites [19] and [66], but debates continue over the biogenicity of these structures [11] and [31]. Because of their uncertain origin, Buick et al. [11] used the term “stromatoloids” to describe these dome-shaped structures. Another important finding at North Pole concerns the oldest evidence of sulfur metabolisms, being either sulfate reducers [47] and/or elemental sulfur disproportionators [37] and [38] (see below).

The top of drillhole PDP1 intersected generally fresh, low-grade vesicular metabasalt of the Maddina Formation from the surface to a depth of 42.72 m (Fig. 3). Rubbly zones at 7.7–8.3 m and 18.1 m, as seen in nearby surface outcrops, are interpreted to represent subaerial basalt flow tops. At the bottom of the metabasalt section, the rocks are recrystallized to a coarser-grained assemblage of chlorite porphyroblasts and anatase (40.28–42.66 m). The lower contact of the metabasalt is a 5 cm-thick zone (42.66–42.72 m) of eutaxitic breccia that lies directly on the Tumbiana Formation. The top of the Tumbiana Formation in the drillhole consists of a 9-cm-thickness of stromatolitic limestone (42.72–42.81 m). This overlies a thin brecciated interval with large clasts of siltstone in a volcanoclastic matrix. Below this (42.89–46.7 m) are interbedded black mudstones (locally pyritic) with local volcaniclastic siltstone (46.3–46.9 m). The interval comprised between 46.7 and 48.53 m is predominantly composed of siltstones displaying cross-bedding, lenticular bedding and flaser structures, with interbeds of laminated mudstones. A stromatolitic build-up is present at 47.02 m, on top of dm-thick interval of fragmented microbial-like sedimentary rocks. Below a brownish oxidized zone between 48.53 and 64.28 m is volcaniclastic siltstone with ripples, and shale (64.28–65.54 m) with some microbial-like laminated intervals between 64.54 and 64.78 m. The underlying interval is dominated by microbial deposits comprising fenestrae-rich laminated sedimentary rocks, finely wavy laminated microbial deposits, pluricentimetric-high domal stromatolites (between 65.65 and 71.3 m; Figs. 2b, 4a) preserving salt pseudomorphs (Fig. 2d), and flat pebble breccias composed of stromatolite fragments (see Fig. 2c). These deposits are interbedded with tuffaceous siltstone (Fig. 2e), mudstone and accretionary lapilli tuff. Locally, individual tuffaceous grains display a filamentous network of ichnofossil-like textures along their boundaries (Fig. 2f). Stromatolite build-ups are again well developed between 76.15 and 79.98 m, with domal structures reaching 20 cm in height. Between 79.98 and 89.56 m are interbedded mudstone, wave-rippled siltstsone (Fig. 4b) and microbialite that grade down into a more volcaniclastic sediment between 86.68 and 89.56 m. The lower part of the hole, from 89.56 to 104.0 m, is monotonous tuffaceous material with accretionary lapilli. Flat-pebble conglomerates are present in the upper 10 cm of the interval and fenestrae occur down to 91.46 m (Fig. 4c). Some microbial mat-like layers occur between 103.31 and 103.85 m. Drilling was stopped at 104.0 m depth.

The drilling site was chosen in the southeastern part of the North Pole Dome, where bedded sedimentary rocks of the Dresser Formation dip 40° south-southeast. Three closely spaced drillholes were sited to the South of the surface outcrops, in the stratigraphic hangingwall, and were oriented at 50°→330° in order to intersect the bedding at right angles (Table 1). The first hole (PDP2a) was drilled with HQ core (75.7 mm diameter) to a depth of 50.6 m through pillowed metabasalt affected by heavy surficial weathering, at which depth the hole was abandoned. No analysis of this core has been made. The second hole (PDP2b) was located further away from the surface outcrops and drilled through metabasalt by RC hammer to a depth of 84.0m, when red flakes of jaspilitic chert emerged from the hole. Diamond drilling with NQ core (47.6 mm diameter) commenced at this depth and continued from 84.0–109.6 m, through the bedded sedimentary rocks, into the underlying metabasalt. The third hole (PDP2c) was sited 14 m further to the south-southeast. Knowing the depth of weathering from the PDP2b hole and the dip and continuity of the bedded sedimentary rocks, the first 69.3 m of this hole was drilled by RC hammer to penetrate below the zone of weathering but remain within the overlying metabasalts. From this depth, drilling continued by NQ diamond drillcore. Metabasalt was intersected to a depth of 92.3 m, where the upper contact of the bedded sedimentary rocks was reached. Diamond drilling was continued through the chert–barite unit to a depth of 114.6 m, well into the underlying komatiitic metabasalt.

The different stratigraphic units in the drill cores are described below from stratigraphic base to top (Fig. 5). The bottom unit consists of hydrothermally-altered komatiitic metabasalt that is transected by numerous veins of grey to black silica and barite. These komatiitic metabasalts preserve relict pyroxene spinifex texture that is now replaced by a carbonate, white mica and pyrite assemblage (Fig. 6a). Vesicles are filled by an early phase of carbonate and a subsequent assemblage of microquartz, recrystallized carbonate rhombs (Fig. 6b) and white mica that was introduced by cross-cutting veins. Locally, the fluid phase that led to the recrystallization of the secondary carbonate is preserved as fluid inclusions surrounding microscopic carbonate grains (Fig. 6c). Cross-cutting veins consist of an early generation of black chert (Fig. 6a) composed of recrystallized carbonate rhombs, abundant carbonaceous material and minor sulfides dispersed in a cherty matrix, and a later generation of grey to white chert with fibrous quartz and carbonate rhombs.

Overlying the komatiitic basalt are 2.5 to 3.6 m of bedded barite composed of coarse crystalline barite fans that alternate with macroscopic pyrite ± sphalerite laminates, and variable amounts of cherty material (Fig. 6d). Centimetre-thick carbonate layers are also present. Locally, large barite crystal fans display sector- and oscillatory-zonation characteristic of hydrothermal environments. Well-preserved overgrowth zones are lined with microscopic, rounded, Ni-bearing sulfides (Fig. 6e; [37]). In most cases, however, barite fans have been extensively recrystallized into a fine-grained assemblage of secondary barite coexisting with fine-grained silica of hydrothermal origin. Associated sulfides lining barite overgrowth zones were also recrystallized into larger, subhedral and Ni-depleted pyrite crystals, which commonly form networks connected to the neighbouring macroscopic sulfide laminates.

In PDP2c, above the bedded barite is a 55 cm thick sequence that is absent in drill core PDP2b. This consists from bottom to top of a hydrothermally-altered volcanoclastic conglomerate (95.0–94.7 m) with rounded clasts of barite and pyrite laminates, a 10 cm thick layer of finely laminated sulfides possibly replacing bedded carbonate (94.7–94.6 cm), a 15 cm thick bed of well sorted, fine-grained felsic volcaniclastic sandstone and a 2 cm thick unit of flat pebble conglomerate composed of curved tabular clasts of bedded felsic ash, now altered to silica and white mica (Fig. 6f).

Variably silicified bedded carbonates overly the felsic volcaniclastic unit in drill core PDP2c and lie directly on the bedded barite in drill core PDP2b. The mineralogy of the bedded carbonate consists of a mixture of ankerite, siderite and calcite, fine-grained silica, pyrite ± carbonaceous material and hematite. Millimetre-to centimetre-scale bedding is well-developed throughout the unit, defined by slight changes in texture and the amount of silica, pyrite, carbonaceous material and hematite (Fig. 7a, c). Although absent in the lower part of the bedded carbonates, the abundance of hematite increases progressively from 84.6 to 84.1 m in PDP2b and from 92.9 to 92.6 m in PDP2c, to become a dominant component of the uppermost 10–30 cm of the drill core (PDP2b, 84.1–84.0 m; PDP2c, 92.6–92.3 m). At several points within the carbonate/jasper beds, there has been a clear infiltration of secondary hydrothermal fluid flow. Hydrothermally-derived silica varies in proportion across the unit, but locally accounts for up to 80% of the rock, where it has replaced the original micritic carbonate and resulted in recrystallization of the micritic primary carbonate to zoned ankerite rhombs with characteristic dusky cores and clear rims. In these layers, it is important to note that recrystallization has also affected the carbonaceous material, which appears remobilized in the replaced matrix (Fig. 7d). This contrasts with the occurrence of micron-thick wavy carbonaceous laminae preserved in the core of large euhedral ankerite crystals in one sample of drill core PDP2b (84.1b; Fig. 7b). In this layer, although the ankerite crystals and carbonaceous lamina are rotated relative to the general bedding plane, the overall trend of the carbonaceous laminate inclusions remains subparallel to the sedimentary bedding. These carbonaceous structures can be interpreted as sub-mm sedimentary layering, potentially the remnants of shallow marine microbial mats preserved in the micritic part of the carbonate rhombs. This morphology is distinctively different from the carbonaceous matter that occurs in adjacent silicified carbonate layers and deeper-seated hydrothermal chert feeder veins. Carbonaceous matter in these rocks appears as dispersed fluffy ‘clots’ remobilized by fluid infiltration without any indication of sedimentary layering (Fig. 7d). Finally, in some of the bedded carbonate samples (e.g. PDP2b sample 84.7), hematite clearly occurs as a secondary pseudomorph of primary ankerite crystals (Fig. 7e). However, not all of the hematite in these sediments is clearly secondary. In several samples (92.5c of PDP2c and 84.6 of PDP2b), some euhedral ankerite crystals contain sub-micron size hematite crystals together with carbonaceous matter (Fig. 7f). This hematite is possibly a primary feature, and was not obviously introduced during later infiltration by oxidized hydrothermal or meteoric waters. Combined carbon and iron stable isotope analysis led to the interpretation that the primary hematite and associated organic material may represent relict activity of anoxygenic photosynthesizers [66].

The upper part of drill core PDP2c (69.3-92.3 m) is composed of amygdaloidal pillow basalt with abundant interpillow hyaloclastite breccia. The mineral assemblage of this unit is a fine-grained mixture of Fe-rich chlorite-carbonate-pyrite ± rutile ± microquartz. Amygdales are filled by a rim of microquartz and a core of coarse-grained carbonate.