Download Sea¯oor hydrothermal alteration at an Archaean mid

J. metamorphic Geol., 2001, 19, 583±599
Sea¯oor hydrothermal alteration at an Archaean mid-ocean ridge
K . K I T A J I M A , 1 S . M A R U Y A M A , 1 S . U T S U N O MI Y A 2 , * AND J . G. LIOU 3
Department of Earth and Planetary Sciences, Tokyo Institute of Technology, Meguro, Tokyo 152±8551, Japan
([email protected])
2
Mineralogical Institute, Graduate School of Science, University of Tokyo, Bunkyo, Tokyo 113±0033, Japan
3
Department of Geological and Environmental Sciences, Stanford University, Stanford, CA 94305, USA
1
ABSTRACT
A hydrothermally metamorphosed/altered greenstone complex capped by bedded cherts exposed in the
North Pole, Pilbara Carton, Western Australia, is interpreted as an accretionary complex. It is distinctive in being characterised by both duplex structure and an oceanic crust stratigraphy. This complex
is shown to represent an Archaean upper oceanic crust with a mid-ocean ridge hydrothermal metamorphism that increases in grade stratigraphically downward. Three mineral zones have been de®ned;
Zone A of the zeolite facies, the prehnite-pumpellyite facies or the lower-greenschist facies at high-XCO2
condition, Zone B of the greenschist facies, and Zone C of the greenschist/amphibolite transition facies.
In Zone A metabasites, Ca-Al silicates including Ca-zeolites, prehnite and pumpellyite are absent
and epidote/clinozoisite is extremely rare. Instead, abundant carbonates are present with chlorite
suggesting high-XCO2 composition in the ¯uid. On the other hand, in Zones B and C metabasites, where
Ca-amphibole+epidote/clinozoisite+chlorite+Ca-Na plagioclase are the dominant assemblages,
carbonate is not identi®ed. The metamorphic conditions boundary of Zones B/C were estimated to be
about 350 uC at a pressure of <0.5 kbar.
Fluid compositions coexisting with Archaean greenstones at the transition between Zones B and C were
estimated by thermodynamic calculation in the CaFMASCH system (T=350±370 uC, P=150±1000 bar)
at XCO2 of 0.012±0.140, such values are higher than present-day vent ¯uids collected near mid-ocean
ridges with low-XCO2 values, up to 0.005. The Archaean seawater depth at the mid-ocean ridge was
estimated to be 1600 m at XCO2=0.06 using a depth-to-boiling point curve for a ¯uid. The carbonation
due to high-XCO2 hydrothermal ¯uids occurred near the ridge-axis before or was coincident with ridge
metamorphism.
Key words: Archaean; greenstone; high-XCO2 ¯uid; hydrothermal metamorphism/ alteration; Pilbara Craton.
INTRODUCTION
Since the discovery of sea¯oor hydrothermal vents at
the Galapagos Spreading Centre in 1977 (Corliss et al.,
1979), a number of workers have investigated the vent
¯uid chemistry and hydrothermal alteration of oceanic
crust at mid-ocean ridges (Campbell et al., 1988;
Butter®eld et al., 1994; Mottl & Wheat, 1994; Alt, 1995;
Von Damm et al., 1995; Charlou et al., 1996). However,
the physico-chemical environment around Archaean
mid-ocean ridges and the features of hydrothermally
recrystallized Archaean oceanic crust have only been
known from the Barberton greenstone belt (3.5±3.1 Ga)
of South Africa (e.g. de Wit et al., 1982; Cloete, 1994;
de Ronde et al., 1994). The geochemical features of
hydrothermal systems at the sea¯oor are well-preserved
in exposed oceanic crust in both Phanerozoic (e.g.
Schiffman et al., 1987) and Archaean (e.g. Ohta
et al., 1996) orogenic belts. Therefore, the study of
*Present address: Nuclear Engineering and Radiological Sciences,
College of Engineering, University of Michigan, MI 48109±2104,
USA.
# Blackwell Science Inc., 0263-4929/01/$15.00
Journal of Metamorphic Geology, Volume 19, Number 5, 2001
parageneses and compositions of secondary minerals
in Archaean oceanic crust is essential to understanding hydrothermal alteration process at Archaean
mid-oceanic ridges.
The goals for the present petrological study are
(1) to determine the compositions and parageneses of
secondary minerals in Archaean oceanic crust; (2) to
estimate P-T-¯uid compositions of hydrothermal
alteration responsible for the observed parageneses;
and (3) to speculate on the characteristics of Archaean
seawater and its interaction with oceanic crust.
The composition of modern seawater is controlled
by two large ¯uxes: river discharges and interaction between seawater and oceanic basalt taking
place mostly within hydrothermal circulation cells at
mid-ocean ridges. The chemistry of seawater is also
controlled through water-vapour circulation between
hydrosphere and atmosphere. In post-Archaean times,
the Earth has had large continental landmasses and
the dominant ¯ux has been continental river discharges
(Windley, 1995). However, Archaean seawater was
predominantly buffered by water/rock interaction at
the sea¯oor (Veizer, 1988). The Archaean atmosphere
583
584
K. KITAJIMA ET AL.
has been considered to have had higher partial CO2
pressure than the modern atmosphere, so Archaean
seawater was enriched in CO2 due to equilibration
with this CO2-rich atmosphere (Owen & Cess, 1979;
Walker et al., 1983; Holland, 1984; Kasting, 1987;
Nisbet, 1995).
One of the best examples for understanding the
Archaean hydrothermal alteration processes and the
physico-chemical environment around mid-ocean
ridge is exposed in the Pilbara Craton, Western
Australia. Awramik et al. (1983) reported wellpreserved ®lamentous-shaped microfossils in chert
as the Earth's oldest fossil from the Pilbara
Craton, Western Australia. They considered that this
microfossil-bearing chert formed in a shallow water
environment, and is representative of a photosynthetic
bacteria, such as cyanobacteria. Maruyama et al.
(1991), however, suggested that the Pilbara Craton,
where these microfossils occur, is an Archaean accretionary complex, and the cherts may have been
formed in a deep-sea environment. Furthermore,
Isozaki et al. (1998) and Ueno et al. (2001) found
some additional localities of fossil bacterium and suggested that the cherts were precipitated in an environment where hydrothermal ¯uid discharged, similar
to the present-day mid-ocean ridges (Isozaki et al.,
1998).
We have selected the North Pole area of the Pilbara
Craton, Western Australia for detailed petrological
study because it has well-preserved, very low-grade
metamorphic assemblages and continuous exposures
of a sequence that is similar to modern ophiolite stratigraphies. The investigated regions are outside the
contact metamorphic in¯uence from tonalite-trondjemite-granodiorite (TTG) plutons within the North
Pole area. We mapped the regions at 1 : 5000 scale and
numerous fresh and altered greenstones, bedded
cherts, bedded and veinlet barites, silica dykes, clastic
rocks and felsic lava samples were collected.
GEOLO GI CAL OU TLIN E
Geology of the North Pole area
The North Pole area of the East Pilbara greenstone
belt is located about 160 km south of Port Headland
and 50 km west of Marble Bar (Fig. 1). The area is
underlain by a granite-greenstone belt of the Warrawoona Group which records a subgreenschist facies
metamorphism (Hickman, 1983). Bedded chert and
greenstone are dominant, with minor silica dykes and
veinlets or layers of barite (Fig. 2). The chert contains
abundant barite and is the lowest sedimentary unit
in the North Pole area (Hickman, 1973). The North
Pole fault-bounded accretionary complex is composed of a series of imbricated piles of pillowed
basalts (>500 m) with minor sheeted dykes and
overlying bedded cherts (>30 m) that locally contain
volcanoclastic or terrigenous sedimentary rocks. These
Archaean sedimentary rocks and greenstones preserve
an oceanic plate stratigraphy similar to those welldocumented sequences of many young circum-Paci®c
accretionary complexes (e.g. the Mino-Tanba belt in
south-west Japan; Matsuda & Isozaki, 1991).
We divided the greenstone in this area into midocean ridge basalt (MORB) and oceanic-island basalt
(OIB) types based on their mode of occurrence and
associated rocks. The MORB type greenstone occurs
with numerous silica dykes and one or two layers of
thick-bedded (several tens of m) cherts on the top of
sequence, and has been subjected to intense carbonation. The OIB type greenstone is intercalated with thin
layers of (=5 m) bedded cherts, and has a few silica
dykes, lacks carbonates, and is composed mainly of
komatiitic basalt.
The greenstone terrane at North Pole consists of
three thrust-fault-bounded greenstone-chert units. The
lowest unit is highly altered with numerous silica dykes
and is described in the next section. The lowest and
middle greenstone units belong to the MORB type,
whereas the uppermost unit is the OIB type. In the
selected area in the lowest unit, the deformation caused
by the adamellite intrusion and its thermal overprint,
which extends only hundreds of metres from the
adamellite (M. Terabayashi, personal communication)
are negligible (Fig. 1).
Geology of the study area
The study area is located to the SE of the North Pole
adamellite which has been dated at 3.46 Ga (Thorpe
et al., 1992) (Fig. 1). The bedded cherts have a total
thickness of up to 70 m, and dip between 30 and
68 u SSE to S. (Fig. 2). The greenstones were divided
into basaltic and doleritic types; most basaltic greenstones are more altered than the doleritic greenstones.
Depending on the degree of alteration, greenstone
varies in colour from pale green to white or light
brown, as secondary minerals such as carbonate, mica
and clay minerals are dominant.
The bedded cherts conformably overlie the basaltic
greenstones and contain no terrigenous debris. Silica
dykes are interpreted to be the fossil pathway of
hydrothermal ¯uids, which have circulated along
normal faults in the upper oceanic crust in the vicinity
of a mid-ocean ridge. The younging direction of this
sequence is indicated by the shape of basaltic pillows
as their tops have convex smooth surface and are
always oriented orthogonal to the strike of bedding
in the cherts. Hence, the chert horizons represent the
original sea¯oor. The Archaean oceanic crust exposed
in this area has an estimated minimum thickness of
up to 1435 m.
Greenstones
Some basaltic greenstones preserve pillow and igneous
intersertal and trachytic textures. Igneous minerals,
however, are completely replaced by secondary
H Y D R OTHERM AL ALT E RATION AT A N A R CH AEA N M OR
5 85
32
32
35
57
38
20
35
76
30
60 80
20
36
30
55
35
45
32
38
30
45
30
58
North Pole
Adamellite
(3.46 Ga)
32
60
34
40
35
26
32
bedded chert
38
no
outcrops
silica dyke
59
42
42
59
25
greenstone
32
basic dyke
41
Fortescue Group < 2.8 Ga
52
36
62
12
40
low-angle reverse fault (thrust)
35
62
55
high-angle normal fault
40
40
40
representative dip & strike of bedding
52
STUDY AREA
36
29
46
30
32
38
45
32
30
40
65
48
30
48
North Pole
68
65
N
42
44
60
44
47
48
52
45
Western
Australia
46
65
50
Pilbara Craton
63
40
40
47
52
45
56
80
5km
60
Fig. 1. Location map of the North Pole area in the Archaean Pilbara Craton (modi®ed after Hickman, 1981) and geological map
of the North Pole area. Detailed ®eld mapping and collection of samples was completed using 1 : 5000 scale topographic maps.
The study area is located in an area 5 km SE of the North Pole Adamellite intrusion.
586
K. KITAJIMA ET AL.
25
middle unit
33
30
47
55 40
42
55
56
40
42
45
64
75
75
30
68
83
42
53
50
56
4
50
80
72
38
70
61
Zone A
Zone A
88
65
42
65
38
36
25 30
Zone B
80
Zone A
(35
60
0˚C
65
)
Zone C
bedded chert
silica dyke
barite
basaltic greenstone
N
doleritic greenstone
Zone C
secondary
minerals
reverse fault (thrust)
high-angle normal fault
chlorite
zone boundary
carbonate
white mica
65
dip & strike of bedding
plagioclase
70
dip & strike of foliation
epidote/
clinozoisite
0
500m
amphibole
Fig. 2. Geological map showing the distribution of secondary mineral assemblages in the study area. Samples were collected where
rocks are exposed and structural data were determined. The poorly exposed areas are shown in the map without structural
elements or secondary minerals.
H Y D R OTHERM AL ALT E RATION AT A N A R CH AEA N M OR
minerals, i.e. plagioclase by albite, chlorite and white
mica, pyroxene by chlorite and white mica, and glass by
quartz, calcite and rutile. Highly altered greenstones
with very ®ne-grained mixtures of quartz and sericite
have no preserved textures. In particular, samples
collected near the silica dykes are extensively silici®ed.
Many basaltic greenstones contain various amounts
of carbonate minerals including calcite, dolomite and
Fe-rich dolomite which often coexist in the same
sample. Carbonates and quartz occur as veins, up to
5 mm in width. Hyaloclastite is observed in only one
locality and is completely replaced by chlorite and
quartz.
The doleritic greenstones are massive and preserve
igneous holocrystalline and intergranular textures.
All igneous minerals are replaced by secondary
phases, except augite, which is partially replaced by
Ca-amphibole along grain margins and cores of crystals. Plagioclase is replaced by albite or oligoclase,
chlorite, white mica and minor epidote/clinozoisite,
and Fe-Ti oxide by titanite, rutile and chlorite. Only
the upper sections of doleritic greenstones contain
minor thin silica dykes.
Chert
In the study area, bedded cherts of various colours
range in thickness from 0.1 to 70 m, and are intercalated with barite layers 0.1±0.5 m thick. Similar
lithologies have been reported from the Barberton
greenstone belt (Paris et al., 1985). Most chert beds
are discontinuous and are terminated by many latestage high-angle normal faults. These faults were
formed due to the adamellite intrusion and have also
cut the thrust planes separating the imbricated piles
of accreted oceanic crust; these faults cut the top of
bedded cherts.
5 87
sulphides, Fe-oxides and ®ne-grained zircon (<3 mm).
Most dyke exposures are black and grey in colour, as
they contain sulphides and oxides; some contain
organic carbon (Ueno et al., 2001).
MINERAL ZONES
Some 300 samples were collected from the area
(see sample locality in Fig. 2); all were thin-sectioned
and some were analysed by EPMA, SEM-EDS, laser
Raman spectroscopy, transmission electron microscopy (TEM) and analytical electron microscopy
(AEM). The greenstones are divided into three mineral
zones (A to C) in an ascending order of increasing
metamorphic grade based on characteristic secondary minerals. The distribution of carbonate, epidote,
chlorite and amphibole is shown in Fig. 2. Most
Zone A greenstones contain carbonate-bearing and
carbonate-free assemblages; only a few samples near
the lower part of this zone contain Ca-Al silicates.
Zone B is de®ned by the occurrence of actinolite
and lacks carbonate-bearing assemblages. Zone C is
the highest metamorphic grade, characterised by the
occurrence of hornblende together with Ca-plagioclase.
The mineral paragenesis of each zone is illustrated
in Fig. 3.
Zone A
This zone is divided into two assemblages: carbonatefree and carbonate-bearing assemblages. Carbonatefree assemblages include:
(A-1) quartz+white mica
(A-2) chlorite+quartz+white mica
(A-3) chlorite+albite+quartz+white mica
(A-4) chlorite+epidote/clinozoisite+quartz+
white mica
Silica dykes
Silica dykes have been termed as `T-chert' (Hickman,
1973); we use the simple term of `silica dyke' rather
than `quartz vein' as some are over 10 m in width
and about 1000 m long and were probably formed by
precipitation of hydrothermal silica along large fractures. More than 1500 silica dykes were recognized in
the North Pole area; 70% are concentrated in the
lowermost unit. The silica dykes cut the basaltic
greenstone and are con®ned to the upper 1000 m of
the greenstone sequences. Silica dykes are capped by,
but do not pass through, the chert. The host greenstones are altered; the degree of alteration is positively
correlated with the number of silica dykes. The silica
dykes decrease in number and in width with increasing
distance from the bedded cherts. They strike SE to
E±W in the upper part of the sequence, and gradually
change to N±S in the lower part with more gentle dips.
The silica dykes are predominantly composed of
®ne-grained quartz and minor Fe, Zn, Pb and Ni
Zone A
Zone B
Zone C
albite~oligoclase
albite~andesine
Epidote/clinozoisite
Actinolite
Hornblende
Plagioclase
albite
Chlorite
Mica
Carbonate
calcite~dolomite
Titanite
Rutile
Quartz
Fig. 3. Schematic mineral parageneses for metabasites from
the study area. Heavy lines show major phases and dashed
lines show minor, or trace phases. The dashed line for rutile in
Zones B and C indicates that it is completely rimmed by
titanite and is not stable.
588
K. KITAJIMA ET AL.
Assemblages (A-1) and (A-2) predominantly occur
in the uppermost part of the greenstones at the contacts with overlying bedded cherts and silica dykes.
These ®ne-grained mineral assemblages were identi®ed either as replacement of basaltic glass or primary
plagioclase and pyroxene, or as mineral aggregates in
veins and amygdules of altered greenstones. Assemblage (A-4) predominantly occurs in the lowermost
section of Zone A near the boundary between Zones
A and B (Fig. 2). Carbonate-free rocks, except (A-4),
are strongly altered and contain chlorite as the
dominant phase.
Carbonate-bearing assemblages were also divided
into four assemblages:
(A-5) carbonate+chlorite+quartz+white mica
(A-6) carbonate+chlorite+albite+quartz
(A-7) carbonate+chlorite+albite+quartz+
white mica
(A-8) carbonate+chlorite+epidote/clinozoisite+
quartz+white mica
All the above assemblages with the exception of
(A-4) also contain minor amounts of very ®ne-grained
rutile. Carbonate-bearing assemblages have similar
modes of occurrence as carbonate-free assemblages.
In Zone A metabasites, carbonate-bearing assemblages
are more abundant than carbonate-free assemblages;
both appear in same domains and are dif®cult to
differentiate.
The predominant carbonates include anhedral calcite and euhedral dolomite rhombs. Some dolomite
rhombs were partially to pervasively altered to hematite. Zone A metabasites contain no index Ca-Al
silicate minerals for determination of the metamorphic
facies and are characterised by quartz+chlorite+white
mica¡carbonate¡albite¡hematite. Trace amounts
of epidote/clinozoisite occur at the boundary between
Zones A and B. Based on the occurrence of albitic
plagioclase, we place Zone A in the lower greenschist
facies or even subgreenschist facies.
Zone B
This zone is characterised by the occurrence of
actinolite and the absence of carbonate. Primary
igneous textures are preserved; except for minor relict
augite rimmed by actinolite+chlorite, all primary
phases are replaced by secondary minerals. Primary
plagioclase is dusted with albite+quartz ¡ white mica.
This zone consists of three mineral assemblages:
(B-1) epidote/clinozoisite+chlorite+actinolite
(B-2) epidote/clinozoisite+chlorite+actinolite+
albite +quartz
(B-3) epidote/clinozoisite+chlorite+actinolite+
albite +quartz+white mica
Albite in this zone contains <10 mol% anorthite
content. The ubiquitous distribution of actinolite,
chlorite, epidote/clinozoisite and albite indicates that
Zone B belongs to the greenschist facies.
Zone C
This zone occupies the middle to stratigraphically
lower part of the greenstones (mainly doleritic
greenstone) and is characterised by two coexisting
Ca-amphiboles (actinolite and hornblende). Primary
igneous textures and minerals are obliterated by
recrystallization. Mineral assemblages of this zone
include:
(C-1) chlorite+albite
(C-2) epidote/clinozoisite+chlorite+albite
(C-3) epidote/clinozoisite+actinolite+albite+chlorite
(C-4) epidote/clinozoisite+actinolite+albite+
oligoclase+chlorite
(C-5) epidote/clinozoisite+actinolite+hornblende+
albite+chlorite
(C-6) epidote/clinozoisite+actinolite+hornblende+
albite+oligoclase+chlorite
All Zone C assemblages contain white mica, quartz
and minor titanite. The lower part of this zone is
occupied by assemblages (C-5) and (C-6). These
assemblages indicate that Zone C belongs to the
greenschist/amphibolite transition facies.
MINERAL CHEMISTRY
Carbonates
Carbonates in Zone A include calcite, dolomite
and Fe-dolomite. Calcite is ubiquitous in Zone A
carbonate-bearing samples. Dolomite becomes predominant at about 500 m from the chert/greenstone
contact and crystals range in size from 0.5 to 2.5 mm.
Some calcite and dolomite occur together as veinlet
minerals that range in size from 0.1 to 0.2 mm across.
Dolomite ranges in Fe content from 0.00 to 1.39 atoms
and has <0.14 Mn atoms per formula unit (p.f.u.),
based on six oxygen (Fig. 4). The Fe/(Fe+Mg) ratio
(XFe) ranges from 0.00 to 0.47, with an average of 0.24.
Fe-rich dolomite occurs dominantly in the upper part
of Zone A.
No apparent trend is shown for compositional
variations of the carbonates with distance from the
chert/greenstone contact, apart from the Mg content
of dolomite, which increases slightly away from the
chert. Calcite at 560 m (96NP322) contains relatively
high Mg with an average value of 0.31 atoms. The Mn
content of calcite is locally higher at 316 m (96NP573)
and 776 m (96NP514) from the chert.
Chlorite
Chlorite is ubiquitous in all zones and replaces pyroxene, plagioclase and interstitial glass. Chlorite compositions are plotted in a Hey's diagram (Hey, 1954)
assuming no ferric iron (Fig. 5). The compositional
H Y D R OTHERM AL ALT E RATION AT A N A R CH AEA N M OR
5 89
Fig. 4. Compositional variation of calcite and dolomite in Zone A greenstones as function of distance from the chert/greenstone
contact.
variation in terms of Al-Fe*-Mg against metamorphic
grade is illustrated in Fig. 6.
Zone A chlorite has a larger compositional variation
and is enriched in Al2O3 compared with that in other
zones; the XFe ranges from 0.35 to 0.80, and the Si
content ranges from 4.96 to 5.97. The XFe and Si
content of Zone B chlorite are in the range 0.23±0.33,
and 5.35±6.43, respectively. The XFe of Zone C chlorite
ranges from 0.23 to 0.71, while the Si content ranges
from 5.09 to 6.89. Chlorite in the lowest grade, Zone
A, has the highest XFe value and lowest Si. Chlorite
in the North Pole basaltic greenstones shows highXFe compositions, because bulk rock composition of
the basaltic greenstone has Fe-rich and Mg-poor
composition compared to the doleritic greenstone.
Chlorite of Zone B doleritic greenstones has Mg-rich
composition.
0.9
Zone A
Zone B
Zone C
0.8
0.7
XFe
Hole 504B
Leg 83
0.6
Horokanai
ophiolite
0.5
Hole 504B
Leg137/140
0.4
0.3
0.2
4.5
5
5.5
6
6.5
7
Si (atoms)
Fig. 5. Chemical compositions of chlorite from Zones A, B
and C plotted on a Hey diagram (Hey, 1954) compared with
those from the DSDP/ODP Hole 504B (Alt et al., 1982;
Laverne et al., 1995) and Horokanai ophiolite (Ishizuka, 1985).
The rock types of Hole 504B Legs 83 are from the pillow/dyke
transition and upper dyke section, and those of Leg 137/140
are from lower dyke section. Zone A chlorite shows a wider
range and higher values in XFe than those in the modern
sea¯oor and Horokanai ophiolite.
White mica
Transparent to pale-yellow mica occurs in all three
zones and replaces the plagioclase and the interstitial
glass of all the observed greenstones. In basaltic
greenstones, it occurs as very ®ne-grained aggregates
(<1 mm). This mica ranges from muscovite or phengite to paragonite in composition. The margarite component is < 0.10, except for sample 99NP112 (j0.21).
The Si content, based on 22 oxygen, ranges from 6.12
to 7.11 atoms p.f.u. and has a mean value of 6.34 in
590
K. KITAJIMA ET AL.
Zone A, and it increases slightly in Zone C with a mean
value of 6.59 (Fig. 7). With increasing distance from
the chert in this area, white mica decreases in AlVI
value at the rate of 0.60 per km, and has increased Fe
and Mg contents indicative of Tschermak substitution
[(Mg,Fe)Si=AlAl]. Zone A white mica has a large
variation in Na/(Na+K) ranging from 0.00 to 1.00;
however, Zone C white mica shows very low values
(<0.30) except for one sample (97NP112), which is
an end-member paragonite that occurs at 1221 m from
the chert.
composition. The variation of pistacite content
[100Fe3+/(Fe3++Al)] is illustrated in Fig. 8, assuming
no ferrous iron and a formula of 12.5 oxygen. The
pistacite content for Zone A samples ranges from 26.1
to 31.0, from 6.5 to 29.8 for Zone B samples, and from
1.9 to 39.80 for Zone C samples. Although scattered,
especially in Zone C, a trend toward an Al-end member
with increasing metamorphic grade is apparent. Some
zoned crystals in Zone C show a decrease in pistacite
content from 27.0 in core to 14.7 in the rim (Fig. 8).
Plagioclase
Epidote/clinozoisite
Epidote/clinozoisite is common in Zones B and C,
but rare in Zone A and has heterogeneous
Plagioclase is common in three zones of the study area.
The primary calcic plagioclase in Zone A is completely
replaced by albite, white mica, carbonates and rare
Al
Zone A
Fig. 6. Chemical compositions of
chlorite from Zones A, B and C plotted
on Al-Fe*-Mg diagram. Fe* means
total iron.
Zone B
Zone C
Fe*
Mg
Zones
0
Distance from chert (m)
A
500
B
1000
C
Zone A
Zone B
Zone C
1500
6
6.2 6.4 6.6 6.8
Si
7 7.2 1.5 2
2.5
3
Al
3.5
(VI)
4 4.5 0
0.5
1
Ca
1.5
2 0
0.5
1
Na
1.5
2
0
0.5
1
K
1.5
2 0
0.2
0.4
0.6
0.8
1
Na/(Na+K)
Fig. 7. Composition-distance variations of analyzed white mica from Zones A (open circles) and C (solid circles) greenstones.
H Y D R OTHERM AL ALT E RATION AT A N A R CH AEA N M OR
5 91
in Zones B and C. The rims of some brown amphibole
are replaced by actinolite and green hornblende. Thus,
these relict amphiboles are excluded from the chemical compositions of the amphiboles in Figs 9 and 10.
In this study, hornblende is de®ned by Si content
(Si<7.20) based on a formula of 23 oxygen.
The Si content of metamorphic amphibole in Zone
B greenstones ranges from 7.26 to 7.92, and 6.18±8.00
in Zone C. The Fe2+/(Fe2++Mg) ratio of amphibole from Zone B ranges from 0.18 to 0.38, with
an average value 0.26. Zone C amphibole shows a
much large Fe2+/(Fe2++Mg) ratio, ranging from 0.11
to 0.91. Based on the Ca-amphibole discrimination
diagram of Leake et al. (1997), Zone B amphibole is
either an actinolite or magnesiohornblende, whereas
those in Zone C are actinolite, magnesiohornblende
and ferrohornblende (Fig. 9).
epidote/clinozoisite. The anorthite content (An%) of
secondary Na-plagioclase reaches 5.6 (albite) in Zone
A and ranges from 0.9 to 14.0 (albite to oligoclase) in
Zone B and 0.2±18.3 (albite to oligoclase) in Zone C
(Fig. 8). The most calcic plagioclase (An%=18.3)
occurs furthest from the chert (1433 m). Plagioclase
in these zones does not show a clear peristerite gap, but
an increase in An% with distance from the chert at the
rate of 4.7 An% per km is apparent (Fig. 8).
Calcic amphibole
Calcic amphibole is common in Zones B and C, and
occurs as a replacement of clinopyroxene. Colourless
to pale green actinolite occurs in Zones B and C. Pale
green to green hornblende occurs only in Zone C. In
addition, minor igneous Ti-rich amphibole is preserved
Zones
0
a
b
Distance from chert (m)
A
500
B
core
1000
rim
core
C
Zone A
Zone B
Zone C
1500
0
Fig. 8. Composition-distance variation
of analyzed epidote/clinozoisite (a) and
plagioclase in (b).
5
10
15
20
25
30
35
40
0
5
10
15
XMg
25
30
35
An%
pistacite content (%)
1.0
0.9
20
tremolite
actinolite
magnesiohornblende
0.5
ferrohornblende
Fig. 9. Chemical compositions of
Ca-amphibole from Zones B and C
greenstones plotted onto the amphibole
classi®cation diagram of Leake et al.
(1997).
Zone B
Zone C
ferroactinolite
0
8
7.5
6.5
Si (atoms)
40
592
K. KITAJIMA ET AL.
Zones
0
Zone B
Zone C
Distance from chert (m)
A
500
B
1000
C
1500
6
6.5
7.0
Si
7.5
8 0
0.2
0.4
Al
0.6
(VI)
0.8
1 0
0.2
0.4
0.6
0.8
1 0
0.2
0.4
XMg
0.6
0.8
1 0
0.1
0.2
Na
0.3
Ti
0.4
0.5 0 0.2 0.4 0.6 0.8 1.0 1.2
3+
Fe
VI
Fig. 10. Composition-distance variation of Ca-amphibole from Zones B and C greenstones. Al , Na and Ti ranges of
Ca-amphibole increase with distance whereas Si and Fe3+ contents decrease with distance.
The Fe3+ content of amphibole was estimated using
the method of Terabayashi (1993) for all examples.
The maximum Na content of amphibole increases
with increasing metamorphic grade due to the edenite
substitution; it ranges from 0.03 to 0.49 in Zone B, and
from 0.01 to 0.81 in Zone C. Amphibole in Zone C
contains higher Ti than those in Zone B; the maximum
Ti content increases with increasing metamorphic
grade.
The amphibole Si content shows no clear compositional gap, although Zone C amphibole shows lower
Si content than Zone B amphibole (Fig. 10). Similar
trends have been reported from other metamorphic
belts such as the Mt. Menzies section and the Elk River
section of the Karmutsen metabasites in Vancouver
Island of Canada (Terabayashi, 1993). The maximum
values of AlVI, Na and Ti increase with distance
from the chert. The Na and Ti contents show maximum
values at 1221 m and 1433 m, respectively. The Fe3+
content decreases with distance from the chert and
shows a maximum value at 672 m in Zone B (Fig. 10).
Rutile, identi®ed by EPMA, EDS and laser Raman,
occurs in all zones with very different grain sizes. Very
®ne-grained TiO2 was analysed by AEM and selected
area electron diffraction (SAED/ Fig. 12a), allowing
it to be positively identi®ed as rutile. In Zone A
greenstones, rutile occurs with ®ne-grained (several
micron) calcite, quartz and minor chlorite (Fig. 12b).
Zones B and C rutile is included by titanite and is not
in direct contact with surrounding minerals such as
Ca-amphibole and epidote/clinozoisite (Fig. 12c),
hence rutile may not be in equilibrium with them.
Small aggregates of rutile in altered greenstone have
been reported from Onverwacht Group (3.5±3.4 Ga),
Barberton greenstone belt, South Africa (Hanor &
Duchac, 1990).
Titanite
DISCUSSION
Titanite is a common mineral in Zones B and C, but it
does not occur in Zone A. It occurs as ®ne-grained
crystals in Zone C basaltic greenstones, and rims rutile
in Zones B and C doleritic greenstones. Titanite compositions are illustrated in terms of Al-Ti/10-Fe in
Fig. 11. Zone B titanite is depleted in Al and Fe and
enriched in Ti, while in Zone C it shows a wide
variation in Al.
Rutile
Timing of the carbonation
Carbonation of greenstone is a common feature in
Archaean greenstone belts. The nature and timing of
carbonation are different in each greenstone belt and
H Y D R OTHERM AL ALT E RATION AT A N A R CH AEA N M OR
provide information about the changes in the physicochemical environment during the alteration process.
Carbonate-bearing assemblages of Archaean greenstones re¯ect the composition of the ancient hydrothermal ¯uid evolved from circulated palaeo-seawater.
The North Pole greenstones have been subjected
to intense carbonation; 40% of collected samples
contain carbonate-bearing assemblages. The timing
Al
basaltic greenstone
doleritic greenstone
5 93
of carbonation of the North Pole greenstones is
estimated by ®eld relations and mineral parageneses.
The carbonate-bearing assemblages are discontinuous
in areal distribution and are bounded by thrust planes
between units and high-angle normal faults due to
adamellite intrusion. Such occurrences suggest that
carbonation occurred on the sea¯oor prior to the
accretion of the greenstone units to the continental
margin. Furthermore, carbonation is recognized in
greenstones adjacent to silica dykes that were formed
during the precipitation of overlying chert near the
mid-ocean ridge-axis (Isozaki et al., 1998). All these
observations suggest that the carbonation of the North
Pole greenstones was a result of alteration with highXCO2 hydrothermal ¯uid and occurred near the ridgeaxis before or coincident with the ridge metamorphism.
Comparison with ophiolites and modern sea¯oor rocks
Ti/10
Fe
Al
Zone B
Zone C
4
1
3
2
Ti/10
Fe
1. prehnite-pumpellyite facies
2. pumpellyite-actinolite facies
3. greenschist to amphibolite facies
4. DSDP/ODP Hole 504B Leg 137/140
Fig. 11. Chemical compositions of titanite from Zones B and
C greenstones compared with those from the prehnitepumpellyite, pumpellyite-actinolite (NystreoÈm, 1983),
greenschist/amphibolite transition facies (Liou & Ernst, 1979)
and 504B Leg 137/140 (Laverne et al., 1995).
We compare the paragenetic sequence of secondary
minerals for the Archaean hydrothermally altered
upper oceanic crust described above with those from
the modern sea¯oor, using examples from the DSDP/
ODP Hole 504B, and some on-land ophiolites. Petrological and metamorphic transitions with depth of
the oceanic crust in these tectonic settings are shown
in Fig. 13. The stratigraphic sequence is similar
among these three settings. The topmost section of
the oceanic crust is capped by pelagic sediments, which
conformably overlie pillowed basalt; coarser-grained
doleritic rocks occur in the deeper parts. In most
ophiolite sequences and the present-day oceanic crust
in the Carlsberg ridge and Mid-Atlantic ridge, gabbroic
rock occurs at about 3 km depth. The overlying
basaltic section grades to doleritic rocks at about
600, 700 and 800 m from the top of the oceanic crust in
Hole 504B, Oman and the North Pole, respectively.
The metamorphic facies of modern oceanic crust
grades from the zeolite and prehnite-pumpellyite facies,
through greenschist, greenschist/amphibolite transition
facies to amphibolite facies with increasing depth.
Similarly, the metamorphic facies for the ophiolite
sequences shown in Fig. 14 change from the zeolite
and prehnite-pumpellyite facies through the greenschist
to amphibolite facies with depth while lower gabbroic
rocks are largely unmetamorphosed. The metamorphic
features observed in the North Pole greenstones are
signi®cantly different from those in modern sea¯oor
and typical ophiolite sections for the upper part of
the sequence. In the North Pole, the lowest grade
metamorphic zone has carbonate-chlorite assemblages
instead of Ca-Al silicates including Ca-zeolites, prehnite and epidote/clinozoisite that are common in
ophiolites and modern sea¯oor rocks (Fig. 13). There
are also differences in compositions of secondary
minerals between North Pole and modern sea¯oor
rocks and ophiolites. The XFe and Si contents in Zones
B and C chlorite overlap those of greenschist or
greenschist/amphibolite transition facies chlorite in the
594
K. KITAJIMA ET AL.
Leg 137/140, ODP/DSDP Hole 504B (Fig. 5). The
differences in the chlorite Al2O3 content of the North
Pole Zones A, B and C cannot be explained by progressive metamorphism with increasing temperature, as
seen in other metamorphic zones (e.g. the Horokanai
ophiolite in the Kamuikotan Zone, Hokkaido, Japan,
Ishizuka, 1985). Zone A chlorite shows Fe-rich and
Si-poor composition as compared with the Hole 504B
Leg 83 chlorite which belongs to subgreenschist facies
(Fig. 5). The large compositional variation of chlorite,
particularly those in Zone A, in part is due to their
occurrence in different domains of carbonate-free or
carbonate-bearing assemblages, and in part is due to
their modes of occurrence as replacement after plagioclase, pyroxene or glass. Furthermore, chlorite composition re¯ects the Fe/Mg ratio of hydrothermal
¯uid and temperature (Saccocia & Seyfried, 1994).
In the study area, Na-rich mica occurs at 1221 m from
the chert. Na-rich micas have been reported from the
Trans-Atlantic Geotraverse (TAG) active hydrothermal mound and are considered to re¯ect high Na/K
ratio of hydrothermal ¯uids (Honnorez et al., 1998).
Epidote/clinozoisite, plagioclase and amphibole compositions show a systematic change with metamorphic
grade. Epidote zoning and variation of pistacite
content indicates that the progressive reaction Fe3+rich epidote+chlorite+actinolite assemblage is temperature dependent and the pistacite content of epidote/
clinozoisite decreases with increasing metamorphic
grade (e.g. the Horokanai ophiolite, Ishizuka, 1985).
The increase of An% in North Pole plagioclase is
consistent with that reported for metamorphic plagioclase from the low-pressure metamorphic facies
series (e.g. Maruyama et al., 1982). The North Pole
titanite compositions overlap with those of East Taiwan
ophiolite, which belongs to the greenschist to amphibolite facies (Liou & Ernst, 1979) and 504B Leg
137/140 (Laverne et al., 1995). Some Zone C titanite
isdepleted in Al and Fe, and enriched in Ti compared to those from the prehnite-pumpellyite facies
and pumpellyite-actinolite facies in Central Sweden
(NystreoÈm, 1983).
Conditions of the Archaean hydrothermal alteration
Fig. 12. (a) A ®ne-grained TiO2 phase identi®ed as rutile by
SAED from [1Å01]. (b) Back-scattered electron image of ®negrained minerals in Zone A greenstone. (c) Back-scattered
electron image of rimmed titanite around rutile in Zone B.
Light coloured needles in lower part of Fig. 12 (c) consist of
titanite and rutile which have replaced exsolution lamella of
ilmenite. The dark coloured part around the needles comprises
quartz and chlorite replacing magnetite.
The boundary between Zones B and C is de®ned by
the occurrence of hornblende together with calcic plagioclase for metabasites, hence it corresponds to the
transition from the greenschist to amphibolite facies.
The metamorphic temperature of this boundary has
been estimated at 350 uC for the Semail ophiolite of
Oman using oxygen isotope systematics (Gregory &
Taylor, 1981), whereas Suzuki (1975) estimated it at
370 uC in a contact metamorphic aureole in Kasuga,
Japan using a calcite-dolomite geothermometer. These
values suggest that the metamorphic temperature
at the boundary of the North Pole greenschist faciesgreenschist/amphibolite transition facies (Zones B/C)
to be about 350±370 uC. This boundary is mapped as
H Y D R OTHERM AL ALT E RATION AT A N A R CH AEA N M OR
5 95
Fig. 13. Simpli®ed lithological and
metamorphic sections showing variation
in metamorphic facies for representative
sea¯oor section and several ophiolite
sequences, and the North Pole, modi®ed
after Ishizuka & Suzuki (1995).
Abbreviations used are: Z: zeolite facies;
PP: prehnite-pumpellyite facies; GS:
greenschist facies; T: greenschist/
amphibolite transition facies; AM:
amphibolite facies; UM:
unmetamorphosed; A: Zone A; B: Zone
B; and C: Zone C.
450
500 bar
Zones B/C
A
SiO2
epidote/
clinozoisite
400
bulk
350
Am
p+E
calc
ite+ p+CO
2+
chl
H
orit
e+q 2O
ua
rt z
T(˚C)
C
bulk
chlorite
amphibole
F
titanite
(2)
(1)
CaO
22
e++CCOO +calcite
tahneitne
stip
rtz
+qua
rutile
GAP
Zone A
SiO2
quartz
bulk
calcite
A
300
TiO2
CaO
rutile
TiO2
epidote/
clinozoisite
bulk
chlorite
calcite
C
250
0
0.1
F
0.2
0.3
0.4
0.5
XCO2
Fig. 14. T-XCO2 section for the CaFMASCH and CaTiSCH
systems at 500 bar. Phase relations were calculated in the
CaFASCH system, with quartz in excess for actinolite,
clinozoisite, daphnite (Fe-chlorite), anorthite and calcite and
for tremolite, clinozoisite, clinochlore (Mg-chlorite), anorthite
and dolomite in the CaMASCH system. Solid line and dashed
lines show equilibria for reactions (1) and (2), respectively. The
shaded area shows Zone A condition at 350±370 uC, and
oblique line area shows Zones B and C conditions. The
ternary diagrams show mineral assemblage and average basalt
composition in the CaTiSCH (SiO2-CaO-TiO2) and
CaFMASCH (ACF) systems.
subparallel to the chert bedding suggesting that the
isothermal lines of sea¯oor metamorphism were subparallel to the sea¯oor and that metamorphic temperature increased with increasing distance from the
chert. This feature is consistent with the modern
sea¯oor metamorphism, hence this area is interpreted
as having preserved primary metamorphism at the
Archaean sea¯oor.
The boundaries of Zones A/B and A/C are not
parallel to the chert bedding, and highly oblique to
the Zones B/C boundary. Zone B assemblages are
restricted to the western part of the study area and are
absent in the eastern part, where Zone A is widespread.
It is to be expected that Zone B should be present in
the eastern part of the area, if the geothermal gradient
and ¯uid composition were similar for both areas.
However, the disparity in the distribution of the zone
boundaries between the western and eastern part can
be explained by a difference in the XCO2 of the ¯uid
that interacted with the greenstones in the two areas.
The change in mineral assemblages from Zone A, with
representative assemblage calcite+chlorite+quartz,
to Zone B, with actinolite+epidote/clinozoisite+
chlorite+albite assemblage, may be related to a reaction in the CaO-FeO-MgO-Al2O3-SiO2-CO2-H2O
(CaFMASCH) system for metabasites. In the presence
of excess quartz, the reaction for the Zones A/B and
A/C boundaries in the eastern part is written as:
3 amphibole+2 clinozoisite+10 CO2+8 H2O (1)
=10 calcite+3 chlorite+21 quartz
On the other hand, Zone A carbonate-bearing
assemblages contain rutile whereas both Zones B and
C assemblages have titanite as an equilibrium Ti-phase,
in addition to other Ca-Al hydrosilicates described in
the previous sections. The difference in Ti-bearing phase
assemblages can be attributed to the following reaction
in the CaO-TiO2-SiO2-CO2-H2O (CaTiSCH) system:
titanite+CO2=rutile+calcite+quartz
(2)
Both reactions (1) and (2) are a function of temperature, pressure and ¯uid composition, and their
temperatures can be estimated by intersection of
these zone boundaries. The boundaries of Zones A/B
and B/C intersect at about 1000 m from the chert.
The hornblende±in boundary of Zones B/C occurs at
a temperature of j350 uC, which can be considered
as an isotherm. Phase relations for reactions (1) and (2)
for the CaFMASCH and CaTiSCH systems were
596
K. KITAJIMA ET AL.
calculated using the THERMOCALC program (Powell
& Holland, 1988); the results are described below.
CaFMASCH system
Solid-solution minerals such as amphibole and chlorite
are assumed to have no Tschermak substitution and
no ferric iron; phase relations were calculated using
Fe-Mg ideal mixing. Non-ideal mixing in CO2-H2O
¯uids was assumed following the approach of Powell
& Holland (1985).
An isobaric T±XCO2 section in the CaFMASCH
system with excess quartz was calculated at 500 bar for
temperatures ranging from 250 to 450 uC (Fig. 14). The
section shows that the Zones B and C assemblage
amphibole+epidote+chlorite is restricted to extremely
low-XCO2 conditions; with increasing temperature,
this assemblage extends towards a higher XCO2 condition. With increase in pressure, reaction (1) shifts
toward more CO2-poor ¯uid compositions.
and C assemblages should be stable in equilibrium
with aqueous-dominated ¯uids.
It appears that phase separation of circulating
hydrothermal ¯uid may have occurred while it rose
to the sea¯oor. Fluid inclusions in barite from same
area show that various liquid/vapour ratios were developed due to boiling or phase separation (Rankin, 1978;
Nijman et al., 1998). This would generate CO2-rich
and H2O-rich ¯uids at shallower depths. In Zone A,
carbonate-free assemblages and carbonate-bearing
assemblages equilibrated with H2O-rich and CO2-rich
¯uid, respectively.
To estimate the ¯uid XCO2 during hydrothermal
alteration, it is necessary to use the pre-separation,
homogenous ¯uid phase, which was estimated using
reaction (1). Reaction (2) could not be used as it yields
a minimum XCO2 of the CO2-rich ¯uid. Two isothermal
P±XCO2 sections (350 & 370 uC) for reaction (1) over a
pressure interval of 150±1000 bar are shown in Fig. 15.
This pressure range equates with the hydrostatic
CaTiSCH system
9000
800
7000
rtar
au
+q
it e
te
rch
lor
cih
t e+
c alc
c te
+ e pi
P (bar)
i b ole
d ote
3000
tz
C
e+
isit
ozo
/clin
400
5000
SEAWATER DEPTH at Archaean MOR (m)
ca
a m ph
600
O
CO2 concentration of Archaean hydrothermal ¯uid
2+
H
2O
At an estimated temperature of 350±370 uC, Zone A
greenstones with the calcite+rutile+quartz assemblage are in equilibrium with high-XCO2 ¯uids, whereas
Zones B and C with titanite and with the assemblage
epidote/clinozoisite+actinolite are associated with
low-XCO2 ¯uids. At 500 bar and 350 uC, there is a
large gap from 0.030 to 0.193 in XCO2 of the ¯uid
(Fig. 14) at the boundaries of Zones A/B and A/C.
Thus the observed disposition of mineral assemblages
in the three zones for the Archaean greenstones from
the study area indicate the in¯uence of ¯uid XCO2 on
the development of secondary mineral parageneses.
A signi®cantly higher XCO2 (e.g. XCO2i0.193 at
350 uC, 500 bar) of hydrothermal ¯uid is necessary
to stabilize calcite-quartz-chlorite assemblages in the
lower part of Zone A greenstones. In contrast, Zones B
370ûC
350˚C
In Zones B and C, titanite is ubiquitous and rims
around relict rutile (Fig. 12c). This difference can be
interpreted by reaction (2). The T±XCO2 relations for
reaction (2) were calculated at P=500 bar (Fig. 14).
At high-XCO2 condition, titanite is unstable with respect
to the assemblage rutile+calcite+quartz at constant
temperatures. This suggests that Zones A, B and C
greenstones located at the same distance from the chert,
at the region of intersection of Zones A/B or A/C and
Zones B/C boundaries, were subjected to nearly
isothermal metamorphism (c. 350 uC). In Zones B
and C, relict rutile rimmed by titanite coexisting
with greenschist and greenschist/amphibolite transition
facies minerals (Fig. 12c). This occurrence may suggest
that these rocks had initially equilibrated with higher
XCO2 ¯uid in an earlier metamorphic/alteration stage.
CaFMASCH system
1000
2
CO
OH2
200
X
1000
0.012 ~ 0.140
0
0
0.05
0.1
0.15
XCO2
Fig. 15. P-XCO2 diagram shows two isothermal sections for
reaction (1) at 350 and 370uC for the CaFMASCH system.
Dotted line shows a depth-to-boiling curve for binary
H2O-CO2 ¯uid at 350uC (Takenouchi & Kennedy, 1964).
Intersection (X) of reaction (1) at 350uC and depth-to-boiling
curve indicates pressure = 260 bar and XCO2=0.06. Arrow
indicates that maximum XCO2 value of present-day sea¯oor
hydrothermal vent ¯uid (XCO2=0.005).
H Y D R OTHERM AL ALT E RATION AT A N A R CH AEA N M OR
5 97
composition by using reaction (1) with a depth-toboiling point curve of H2O±CO2 ¯uid shown in Fig. 15
(Takenouchi & Kennedy, 1964). At 350 uC these two
reactions intersect at P=260 bar and XCO2=0.06
suggesting that the North Pole seawater depth was
about 1600 m and is similar to modern mid-ocean
ridges.
pressure (seawater depth) developed at modern midocean ridges. At the interpreted condition (Ptotal=
P¯uid=150±1000 bar and T=350±370 uC), reaction
(1) buffers the ¯uid XCO2 in the range 0.012±0.140.
The maximum XCO2 value of modern mid-ocean
ridge hydrothermal vent ¯uid is 0.005; subgreenschist
facies assemblages are common in drilled and dredged
metabasalt samples (Alt et al., 1986; Von Damm et al.,
1985). This value cannot be compared with those
deduced from the present study for the Archaean
hydrothermal ¯uid directly, because they were estimated at different P-T conditions. However, the hydrothermal ¯uid for the North Pole greenstones is
estimated to have 12 times higher CO2 concentration
than present-day hydrothermal ¯uid. Hence, Ca-Al silicates such as Ca-zeolite, prehnite, epidote/clinozoisite
and titanite are not stable in Archaean upper oceanic
crust.
If Archaean ¯uid ranges in XCO2 from 0.012 to 0.140
and has undergone subsea¯oor phase separation or
boiling at 1 km depth from sea¯oor, the pressure range
for this depth is estimated to be 180±370 bar (Fig. 15).
This pressure range requires an overlying minimum
water column of 1800 m, represented by a seawater
depth of 800 m and a water column from the sea¯oor
to the depth of phase separation/boiling at 1000 m.
Furthermore, we constrain the pressure and ¯uid
CONCLUSIONS
A schematic model for the Archaean oceanic crust
consisting of gabbro dolerite dyke sequence and pillow
basalt capped by bedded chert is shown in Fig. 16.
Such layered oceanic crust created at the spreading axis
has been subjected to ridge-hydrothermal alteration
similar to the modern ocean crust. The shallow transition depth for the greenschist to amphibolite facies
suggests a higher geothermal gradient for Archaean
ocean-ridges (see Fig. 13). Figure 16(b) shows a zonal
distribution of hydrothermal metamorphism/alteration
of the Archaean oceanic crust. The formation of a
heterogeneous ¯uid as a result of phase separation of
an ascending homogeneous ¯uid at depths of about
1000 m and 350 uC, plus the disposition of Zones A, B
and C assemblages with a few representative silica
dykes formed along fractures for hydrothermal vents
are also illustrated. The 3.5 Ga greenstone sequence
(b)
precipitation
0
Archaean Mid-Ocean Ridge
sea level
RIDGE AXIS
seawater depth
1600 m
(b)
bedded chert
magma chamber
id
c
la
500
1000
Zone A
phase separation
or
boiling
Zone A
PHASE SEPARATION
or
BOILING
Zone B
Zone B
350˚C
Low XCO2
stable assemblage and mineral
BASALT
DOLERITE
calcite + rutile + quartz
Zone C
u
fl
Zone A
H2O-rich fluid
and
CO2-rich fluid
(XCO2 > 0.06)
homogeneous
fluid
n
u
c
ir
CRUST
High XCO2
stable assemblages
calcite + chlorite + quartz
o
ti
OCEANIC
Distance from seafloor (m)
(a)
bedded chert
hert
NATURE OF FLUID
heterogeneous
fluid
hydrothermal vents
1500
low-XCO2 fluid
(XCO2 < 0.06)
amphibole + epidote/clinozoisite
Zone C
titanite
GABBRO
Fig. 16. Schematic cross-section of the suggested Archaean mid-ocean ridge. (a) Hydrothermal ¯uid circulates in the upper oceanic
crust which consists of bedded chert, basalt, dolerite and gabbro. (b) Enlarged view showing detailed cross section of Archaean
uppermost oceanic crust with respect to the nature of hydrothermal ¯uid. Circulating ¯uid rises along normal faults and fractures,
and is subjected to phase separation or boiling at 350 uC and roughly 1000 m distant from the chert. Phase separation or boiling
resulted in H2O-rich and CO2-rich ¯uids in the upper part, where high-XCO2 assemblages are stable. Thick-bedded chert
precipitated from vented hydrothermal ¯uid at the top of the greenstone sequence.
598
K. KITAJIMA ET AL.
from the North Pole Pilbara Craton, Western Australia
preserves Archaean mid-ocean ridge hydrothermal
alteration. The upper 1000 m of the greenstone was
subjected to carbonation by high-XCO2 hydrothermal
¯uid with an estimated XCO2 of 0.012±0.140. Ca-Al
silicates, common in present-day oceanic crust and
ophiolites, do not occur in the Archaean upper oceanic
crust.
During the ascending path, a high-XCO2 ¯uid may
have separated into H2O- and CO2-rich phases at
depths of about 1000 m below the sea¯oor. Interaction
of upper oceanic crust with discharged H2O ¯uid
would result in chlorite-white mica-quartz assemblages
(carbonate-free assemblages of Zone A), whereas
those with CO2-rich ¯uid develop carbonate-bearing
assemblages. Because of the dynamic processes with
continuous change in temperature as well as ¯uid
composition during the migration of the oceanic crust
away from the spreading ridge, the parageneses and
compositions of secondary minerals in both basaltic
and doleritic sections are extremely complicated.
The schematic model for the Archaean mid-ocean
ridge shown in Fig. 16 is provisional. Nevertheless,
our observations for metamorphism/alteration of the
Archaean oceanic crust indicate that the Archaean
hydrothermal ¯uid was higher in XCO2 compared to
its modern analogue. The North Pole carbonation
occurred near the ridge-axis before or coincident with
ridge metamorphism.
ACKNOWLEDGEMENT S
We thank Y. Isozaki, M. Terabayashi, Y. Kato, T.
Kabashima, Y. Ueno and K. Nakamura for assistance
during the ®eldwork. Collaboration in the ®eld with
A. Thorne and A. H. Hickman was helpful. This
research was supported by the Ministry of Culture and
Education of Japan (International Scienti®c Research
Program; Field research nos. 06041038 and 08041102
and Intensi®ed Study Area Program, no. 259 1995±97).
Preparation of the manuscript was completed while
Liou was taking sabbatical supported by the Tokyo
Institute of Technology and the NSF Center of
Global Partnership Fellowship, NSF INT-9820171.
We thank Drs C. E. J. de Ronde, J. C. Alt and
D. Robinson for critical reviews that materially
improved the manuscript.
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Received 26 July 2000; revision accepted 26 March 2001.