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Fracture sealing and fluid overpressures in limestones ofthe Jabal Akhdar dome, Oman mountains
C. HILGERS1, D. L. KIRSCHNER2, J . -P. BRETON3 AND J. L. URAI1
1Geologie-Endogene Dynamik, RWTH Aachen, Aachen, Germany; 2Department of Earth and Atmospheric Science, Saint
Louis University, St Louis, MO, USA; 3BRGM Oman Branch, Ruwi, Oman
ABSTRACT
Fractures are important conduits for fluid flow in the Earth’s crust. To better understand the spatial and tem-
poral relations among fracturing, fracture sealing, and fluid flow, we have studied fractures, faults, and veins
in a large dome (Jabal Akhdar) in the Oman mountains. Our work combines the results of meso- and micro-
structural analyses and stable isotope analyses. Seven generations of fractures and veins have been identified
in the carbonate-dominated dome. The earliest generations of veins developed during extension and subsi-
dence of the Mesozoic basin. These veins formed in the inclined segments of bedding-parallel stylolites and
in extensional fractures that are perpendicular to bedding (#1 and #2, respectively). These extension-related
veins are truncated by bedding-parallel veins (#4) that formed during top-to-north bedding-parallel shear of
both the northern and southern limbs of the dome. These veins are consistent with a change in stress regime
and may be related to an earlier generation of strongly deformed pinch-and-swell veins (#3) that are exposed
locally on the southern limb of the dome. Normal faults contain a set of en-echelon tension gashes (#5) and
veins emplaced in dilational jogs along the fault planes (#6). In the northern part of the dome, veins (#7)
associated with thrusts post-date the normal faults. Samples of veins and their host rocks were analyzed
to provide information on fluid-rock interaction in the dome and the scale(s) of fluid movement. Oxygen iso-
tope values range from +16.2 to +29.3&; carbon isotope values range from 0 to +3.6&. The results of the
structural and isotopic analyses are consistent with the early veins (#2–#5) having precipitated from overpres-
sured fluid in a isotopically rock-buffered system. During normal faulting (#5 and #6), a more open system
allowed external fluid to infiltrate the dome at drained conditions and precipitate the youngest sets of veins
(#6 and #7).
Key words: paleo-stresses, stable isotopy, veins
Received 25 February 2005; accepted 17 January 2006
Corresponding author: Christoph Hilgers, Geologie-Endogene Dynamik, RWTH Aachen, Lochnerstr. 4–20, D-52056
Aachen, Germany.
Email: c.hilgers@ged.rwth-aachen.de. Tel: + 49 241 809 5723. Fax: + 49 241 809 2358.
Geofluids (2006) 6, 168–184
INTRODUCTION
Fluid overpressures are a well-known phenomena in many
sedimentary basins (e.g. Ortoleva 1994; Bjørlykke 1997;
Law et al. 1998). They cause a reduction of effective stress
and enhance the potential of rock failure (Hubbert &
Rubey 1959). Fluid overpressures may cause hydraulic frac-
turing, which increases the bulk permeability of rock and
may result in connected fracture networks (Cox et al.
2001). This increases significantly flow rates (Taylor
1999), and allows the influx of external fluids.
Many veins have been shown to grow from supersaturat-
ed solutions in hydraulic extension fractures, and thus rep-
resent paleo-fluid conduits that formed in an overpressured
environment. Thus, vein formation can involve both frac-
turing and sealing at elevated fluid pressures (Ramsay
1980). Sealing may be caused by (i) a pressure drop and
associated secondary effects such as CO2 partial pressure
reduction or boiling, and (ii) an increase of pore sizes such
as during the formation of a fracture, which enhances the
nucleation potential and reduces the equilibrium concen-
tration of the fluid (e.g. Putnis et al. 1995; Barnes & Rose
Geofluids (2006) 6, 168–184 doi:10.1111/j.1468-8123.2006.00141.x
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Journal compilation � 2006 Blackwell Publishing Ltd
1998; Putnis & Mauthe 2001). Fracture sealing from
externally derived fluids may also be driven by (iii) a tem-
perature gradient between fluid and wall rock, or (iv) a dif-
ferent wall rock composition exposing the fluid to different
Eh-pH conditions (e.g. Skinner 1997; Robb 2005, pp.
154–159).
The deformation and fluid-related histories of the lime-
stones can be deciphered from the numerous crosscutting
faults, fractures, and veins in the carbonates. To document
the deformation and fluid-related histories, we combine
the observations and analyses of our field study with micro-
structural observations of the veins, and stable isotope ana-
lyses of host rocks and veins to derive the deformation and
fluid histories of the carbonates during deformation.
GEOLOGICAL SETTING
The Oman mountains of the Alpine-Himalayan chain
formed during northeast-directed subduction of Arabia
below the Eurasian plate (Fig. 1). Intra-oceanic subduction
started in the Cenomanian and continued into the Middle
Turonian to early Campanian with the obduction of the
Samail ophiolite (Boudier et al. 1985; Beurrier et al. 1989;
Hacker 1994; Hacker et al. 1996; Breton et al. 2004 and
references therein). Several tectonic windows below the
ophiolite are present in the region including the Jabal Akh-
dar, which is the focus of this study.
Metamorphic facies associated with subduction range
from upper anchizone in the Jabal Akhdar to blueschist
and eclogite facies in the Saih Hatat tectonic window (le
Metour 1988; Breton et al. 2004; Searle et al. 2004). Peak
HP-LT metamorphism in the Saih Hatat is dated at about
80 Ma, concurrent with subduction (Warren et al. 2003),
while other data are consistent with peak metamorphism at
110 Ma and post-100 Ma exhumation (Gray et al.
2004a,b) (Fig. 1). Locally, in the Saih Hatat area, the
metamorphic gradient is more complex because of interca-
lation of different rock packages by numerous shear zones
and to structural thickening (Gregory et al. 1998; Miller
et al. 2002).
The Jabal Akhdar dome, which is located in the central
part of the Oman mountains, is a large 2500 km2 box-
shaped dome with an amplitude of 3 km and a wavelength
of 70 km (Searle 1985) (Fig. 2). Exposed in the core of
the Jabal Akhdar dome are pre-Permian rocks that are
unconformably overlain by Middle Permian to Cenoma-
nian carbonates (Glennie et al. 1974). This 2.5-km-thick
sequence of carbonates was deposited on the subsiding
southern passive margin of the Tethyan ocean (Hanna
1990; Mann et al. 1990; Pratt & Smewing 1993; Masse
et al. 1997, 1998; Hillgartner et al. 2003). The carbonates
are very well exposed laterally and vertically, with excellent
1-km-high exposures in numerous wadis that transect the
dome.
To date, we have focused our study on the Cretaceous
limestones and clayey limestones that are exposed in the
Wadi Mistal and Wadi Bani Awf of the north limb and
Wadi Nakhar of the southern limb (Fig. 2). These platform
500 km
N
Zagros suture
Zagros fold belt
Riyadh
Kuwait
Muscat
TehranBaghdad
ARABIAN PLATE
AFRICAN PLATE
EURASIAN PLATE
Oman mts.
10°
20°
40°
10°
20°
30°
40°60°50°40°
60°50°40°
Maastrichtian-tertiary sediments
Samail ophioliteHawasina allochthonous sediments
Autochthonous permian-cretaceous platform carbonates Lower units incl. eclogitesPre-permian basement
Today’s stress orientationThrusts
50 km
N
Ibra
Muscat
Nizwa
Saih Hatat window
Jabal Akdhar window
Hawasina window
Gulf of Oman
Fig. 1. Three tectonic windows are present in
the Oman mountains. The Jabal Akhdar tectonic
window is a dome with Infracambrian sedimen-
tary rocks in its core. The Hawasina nappes and
the Samail ophiolite were obducted southward
on autochthonous basement and Mesozoic plat-
form carbonates (modified after Glennie et al.
1973; Hanna 1990; Miller et al. 2002). (inset)
Arrows indicate orientations of recent stress field
in regional map. The Zagros mountains and
Makran subduction zones are the suture zone
between Eurasian and Arabian plates (modified
after Konert et al. 2001; Breton et al. 2004; str-
ess orientations from http://www.world-stress-
map.org).
Fracture sealing and overpressures in limestone 169
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Journal compilation � 2006 Blackwell Publishing Ltd, Geofluids, 6, 168–184
carbonates, which transition into deep-water facies north-
east of the dome (Masse et al. 1997; Hillgartner et al.
2003), have been described in detail by Breton et al.
(2004).
The uppermost autochthonous carbonate unit is com-
posed of Turonian to Santonian megabreccias and olisto-
liths of the Muti formation. These sediments were derived
from an outer carbonate platform to the northeast and
document the transition from a passive continental margin
to a foredeep basin (Robertson 1987; Breton et al. 2004).
The pelagic matrix of the megabreccias is Middle to Late
Turonian in the northern part of Jabal Akhdar and Conia-
cian-Santonian in the southwest part of Jabal Akhdar (le
Metour 1988; Rabu 1987; Breton et al. 2004), consistent
with migration of the depocenter toward the south during
subduction (Robertson 1987; Breton et al. 2004).
The Muti formation is unconformably overlain by allo-
chthonous Permian-Cretaceous volcano-sedimentary rocks
of the Hawasina nappes and the Samail ophiolite. The
Samail ophiolite comprises two magmatic sequences. The
first one is dated Albian to early Cenomanian (Beurrier
1988) and formed along a 500-km-long paleo-ridge (Nico-
las et al. 1988). The second sequence is dated at 97–
94 Ma (Tilton et al. 1981; Beurrier 1988) and is related to
the intra-oceanic subduction. Obduction of the Samail
ophiolite onto the continental margin north of Jabal Akh-
dar started in the Middle Turonian. The Hawasina and Sa-
mail tectonic pile was then thrust southward over the Jabal
Akhdar area in Late Santonian (Bechennec et al. 1988;
Breton et al. 2004). Almost 450 km of displacement of
the Semail ophiolite obduction occurred between 95 and
80 Ma (Warburton et al. 1990; Breton et al. 2004). Shear
indicators at the base of the ophiolite are consistent with
top-to-south shearing over the Jabal Akhdar dome (Bou-
dier et al. 1988).
The thrust stack is unconformably overlain by autochth-
onous Maastrichtian and early Tertiary limestones, which
were deposited after nappe emplacement (Glennie et al.
1973; Hanna 1990). The distribution and onlap of Ter-
tiary sediments occurred during Eocene to Miocene uplift
of Jabal Akhdar (Searle et al. 2004), which resulted in the
erosion of the overlying Hawasina and ophiolite nappes in
the central part of Jabal Akhdar. The Jabal Akhdar dome
formed during multi-phase deformation from late Creta-
ceous to Early Paleocene and Miocene-Pliocene (Glennie
et al. 1974; Searle 1985; Glennie 1995; Poupeau et al.
1998; Gray et al. 2000). During this time, the dome prob-
ably experienced maximum temperatures of approximately
200�C and maximum overburden pressures of 200–
400 MPa that increased from SW to NE across the dome
(le Metour et al. 1990; Breton et al. 2004).
VEINS
We focused our study on the ubiquitous veins that outcrop
in the dome in order to understand the dome’s structural
Fig. 2. Geologic map of Jabal Akhdar dome. The outer limb consists of autochthonous Mesozoic carbonates, which were sampled in Wadi Bani Awf and
Wadi Mistal on the northern flank, and Wadi Nakhar on the southern flank (modified after Bechennec et al. 1993).
170 C. HILGERS et al.
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history and paleohydrology during formation and exhuma-
tion. We identified seven generations of calcite-rich veins
in three wadis of the Jabal Akhdar dome. The veins and
overprinting relations are similar in the three wadis, and
are thus discussed together. We used the veins’ relations to
bedding surfaces, stylolites, boudins, and normal faults to
determine their relative ages (Table 1).
Stylolite veins (#1)
The first generation of calcite veins formed in steeply
inclined segments of stylolite seams. The stylolite seams are
parallel to bedding, enriched in clay, are up to 1 cm thick,
have irregular shape indentations, and would be described
as composite seams according to the nomenclature of
Guzzetta 1984). In Wadi Nakhar, these stylolites have
accommodated up to 10 cm of shortening by dissolution
(Fig. 3A).
Bedding-normal veins (#2)
A set of vertical calcite veins crosscut the stylolites, are per-
pendicular to bedding (Fig. 3A), and are truncated by nor-
mal faults. On the north limb of the dome, bedding-
normal veins form in the necks of boudin layers (Fig. 3B).
These boudins are absent in Wadi Nakhar, where the same
vein generation is present in undeformed beds. This
north–south decrease in deformation intensity is consistent
with the decrease in metamorphic gradient and overburden
toward the south. The veins strike north–south to north-
west–southeast, parallel to the veins in the stylolite seams.
Pinch-and-swell veins (#3)
On the southern limb of Jabal Akhdar, veins are locally
stretched to form pinch-and-swell structures in cleaved
marls (Fig. 4A). The spaced cleavage dips toward the west
and contains small pressure shadows around pyrites.
Grooves on the pinch-and-swell vein surface are consistent
with top-to-east shear. Staining in the laboratory of the
pinch-and-swell veins has revealed several stages of frac-
ture-sealing events. Calcite breccia within a quartz vein
indicates that the vein was first sealed with calcite, brecciat-
ed, and then resealed by a second generation of calcite
(Fig. 4B,C).
Crosscutting relations between pinch-and-swell veins
and other veins have not been observed, thus relative tim-
ing between these vein sets is difficult to establish. How-
ever, the pinch-and-swell veins are truncated by and thus
pre-date brittle normal faults. These veins may be associ-
ated with the bedding-parallel veins (#4).
Bedding-parallel veins (#4)
Two types of bedding-parallel veins have been observed.
Thin (<5 cm thick) layered bedding-parallel veins (#4.1)
extend for several tens of meters, truncate subvertical veins
of set #2 (Fig. 3C), and are offset locally by en echelon
vein arrays and associated normal faults (#5 and #6)
(Fig. 3D,E). The veins contain host rock inclusion bands
that are arranged sub-parallel to the vein wall. Locally, the
vein is folded with an axial plane dipping toward the south,
consistent with top-to-north shearing. The second type of
bedding-parallel vein formed in dilational jogs (#4.2) that
opened several centimeters wide during bedding-parallel
shear (Fig. 4D). On both limbs, this second type of bed-
ding-parallel vein formed during top-to-NNE shearing. A
relative chronology between these bedding-parallel veins
has not been established by crosscutting relationships.
Normal faults and associated veins (#5 and #6)
Normal faults with throws up to several hundreds of
meters are associated with drag folds, slickensides, grooves,
and two vein sets on both sides of the dome (Fig. 3D,E).
The dips of conjugate normal faults change across the Jabal
Akhdar and are consistent with faulting having occurred
prior to dome formation. The first set of veins (#5) is com-
Table 1 Overview of vein sets and evidence for relative time relationship.
Vein
generation Description Evidence for relative age relationship Regime
#1 Stylolite vein Cross cut by extension veins (#2) r1 normal to bedding
#2 Extension vein, boudinage vein Truncated by bedding-parallel vein (#4) r1 normal to bedding
#3 Pinch & swell vein Undeformed extension vein nearby, truncated by normal fault (#5, 6),
relative timing to #4 unclear
r1 oblique to bedding
#4 Bedding-parallel vein Displaced by normal faults (#5), top-to-the north directed shear indicators
indicating bedding parallel slip, crosscutting relationship of #4.1 and #4.2 unclear
r1 oblique to bedding
#5 En echelon extension vein arrays Aligned along and often displaced by normal faults (#6), no visible offset by
bedding parallel slip (i.e. #4)
r1 normal to bedding
#6 Veins in dilation sites of normal faults Truncate en echelon veins (#5) r1 normal to bedding
#7 Thrust veins Cross cuts #2, #4, and #5 r1 oblique to bedding
Fracture sealing and overpressures in limestone 171
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Journal compilation � 2006 Blackwell Publishing Ltd, Geofluids, 6, 168–184
posed of en-echelon arrays of veins that are oriented per-
pendicular to bedding and strike parallel to the normal
faults. The en-echelon veins are offset by normal faults and
bedding planes that slipped during drag folding
(Fig. 4E,F). The second vein set (#6) occurs in dilational
jogs along the normal faults.
In contrast to the earlier vein sets #1–#5, vein set #6
often does not seal the fault completely and shows euhe-
dral crystals grown in open cavities. The surfaces of these
veins are striated sub-vertically and sub-horizontally. Cross-
cutting relations are consistent with sub-horizontal move-
ment having occurred after sub-vertical movement.
15 cm
15 cm 2 m
3 cm 1 m
30 cm
#4
#4#5
#6
#1
#2
#4
#2
#7
#6
#4
#2
A B
C D
E F
Fig. 3. (A) Bedding-parallel stylolites contain calcite veins (set #1) in their steep limbs of the teeth (arrow). Some stylolites are cross-cut by vertical calcite
veins. Wadi Nakhar. (B) Extensional veins (#2) formed in boudin necks in the northern limb of Jabal Akhdar (indicated by arrows). Veins are normal to bed-
ding. (C) Thin dark rock slivers within a horizontal vein (#4) crosscut a vertical vein (#2). Wadi Nakhar. (D) Normal faults displace two horizontal veins (#4)
(indicated by white arrows). The normal fault on the left contains a thick vein that formed in a dilational jog of the fault (#6). Wadi Nakhar. (E) En echelon
tension gashes (#5) are associated with normal faulting. Slip of the normal fault causes dragged (folded) adjacent limestone beds. Rectangular area is
enlarged in Fig. 4. Wadi Nakhar. (F) Reverse faults and thrusts (#7) displace earlier normal faults and boudins (#2), and are associated with fault-bend fold-
ing. Arrows denote two limestone beds with regularly spaced boudin veins, that are oriented normal to bedding. Wadi Mistal.
172 C. HILGERS et al.
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Veins in late thrust faults (#7)
Small thrusts, associated folds, and reverse faults are pre-
sent in the northern limb of the dome at Wadi Mistal
(Fig. 3F). They dip WSW with slip toward the northeast.
The thrusts, which offset the normal faults, contain blocky
calcite veins on the primary thrust surfaces. These late
thrusts may have formed during one phase of doming of
Jabal Akhdar in the Tertiary (Glennie et al. 1974; Gray
et al. 2000).
Pencil
3 cm
30 cm
3 cm
S0
S N
S N
#4.1#5
#6
A
E
C
B
D
#3
5 cm
SN
#4.2
2 cm
Fig. 4. (A) Pinch-and-swell veins are oriented oblique to bedding and cleavage and indicate intense stretching. They dip sub-parallel to cleavage (bedding is
horizontal). Wadi Nakhar. (B) Overview of stained pinch-and-swell vein (#3). Note angular quartz clasts (light color) within the calcite vein (dark color).
Quartz fragments are located in both the pinch-and-swell regions, indicating that the quartz vein filled the whole vein prior to stretching. Wadi Nakhar. (C)
Stained vein sample from the pinch-and-swell structure. Calcite (dark color) is located along the vein–host rock interface. Central part of vein is filled with
quartz (light color). Locally, angular fragments of calcite ‘float’ within quartz vein consistent with quartz precipitation after calcite. (D) Bedding-parallel calcite
veins form in dilational jogs (#4.2) and displace some bedding-normal veins (#2). Shear criteria show top-to-the NNE shear sense. Wadi Nakhar. (E) Detail of
Fig. 3E. Tension gashes (#5) crosscut bedding-parallel veins and are displaced along the bedding plane close to the normal fault. Steeper parts along the nor-
mal faults form dilation jogs (#6) and are filled with blocky calcite veins. The normal fault off-sets bedding-parallel veins (#4.1). Wadi Nakhar.
Fracture sealing and overpressures in limestone 173
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METHODOLOGY OF ISOTOPE ANALYSES
Stable isotope analyses of host rocks and veins were made
to determine the history of fluid-rock interaction and fluid
flow during the deformation history. Samples were collec-
ted from approximately 50 outcrops primarily in the three
wadis of the Jabal Akhdar dome. The samples were slabbed
with a water-lubricated saw and micro-drilled. The result-
ing powders were then analyzed on an automated carbon-
ate-reaction device that is connected to a continuous-flow,
gas-source, isotopic ratio mass spectrometer at Saint Louis
University. Approximately 0.5 mg of each powder was
reacted with H3PO4 at 90�C for several hours prior to ana-
lysis. Approximately one in-house standard was analyzed
for every five unknowns. The in-house standard’s d13C
value of )2.33& and d18O of 24.72& were calibrated to
NBS-19 standard values of 1.95 and 28.64&, respectively.
The standard deviations of the in-house standard in indi-
vidual sets of analyses (generally 30 analyses for each set)
were between 0.01 and 0.08& for d13C and 0.08 and
0.14& for d18O (1r; n ¼ 58). Some unknown powders
were analyzed two to four times; the reported values are
averages of these analyses.
RESULTS OF ISOTOPIC ANALYSES
Host rocks have d13C values between 0.5 and 3.6& and
d18O values between 23.4 and 28.4& (Table 2, Fig. 5).
Similar values have been obtained from Aptian to Cenoma-
nian age carbonates that were deposited in the paleo-Tet-
hys sea and similar age oceans (cf. Scholle & Arthur 1980;
Renard 1986; Burkhard & Kerrich 1988; Weissert et al.
1998; Jenkyns & Wilson 1999; Ghisetti et al. 2001; Skel-
ton 2003, p. 169, 267).
Veins have d13C values between 0.1 and 3.6& and d18O
values between 16.2 and 29.3& (Fig. 5). Most of the early
veins (#1–5) have d13C and d18O values similar to their
adjacent host rock values (Fig. 6), while some veins of gen-
erations #6 and #7 have d18O values that differ signifi-
cantly from their host rock values. The isotopic values vary
slightly within individual veins (Fig. 7), though much less
than the range in values among vein generations #1–#7
(Fig. 8). The transition from a more closed to open fluid
system can be seen in the increased variability of the d18O
data that occurs from vein sets #5 and #6 (Fig. 8).
Some insights into the potential sources of fluids
involved in vein formation can be gained by looking at the
isotopic values of modern precipitation, groundwater, and
carbonate cements formed in surface alluvium, travertine,
and aeolianite deposits from northern Oman. Modern pre-
cipitation in the Oman region of the Arabian Peninsula has
mean annual d18O values of approximately )4 to 0&
SMOW (Rozanski et al. 1993; Clark & Fritz 1997, p. 51),
Table 2 Stable isotope data of host rocks and veins. The vein generations
were classified according to the relative crosscutting relations.
Sample no. Type (generation) d18O (VSMOW) d13C (PDB)
Wadi Mistal
1(a) Vein (#6) 25.47 3.38
2 Vein (#6) 26.20 3.36
2 Host rock 25.89 3.57
3 Host rock 26.53 2.67
4(1) Vein (#2) 25.76 3.51
4 Vein (#2) 26.04 3.61
4 Host rock 25.74 3.41
5(2) Vein (#7) 25.66 3.59
5(4) Vein (#7) 25.50 3.43
5(3) Vein (#7) 25.73 3.20
5(5) Vein (#7) 24.18 3.04
5 Vein (#7) 25.31 3.33
5 Host rock 23.44 2.01
5 Host rock 24.12 1.36
6 Vein (#4) 26.14 3.53
6 Host rock 25.51 3.50
8(a) Vein (#5) 25.48 3.43
8(b) Vein (#5) 25.79 3.51
8(c) Host rock 25.25 3.03
8(c) Vein (#5) 25.46 3.53
8(d) Vein (#2) 26.07 3.45
Wadi Bani Awf
9(a) Vein (#7) 16.23 1.66
9(b) Vein (#7) 19.38 1.87
9(b) Host rock 27.55 2.22
9(a) Host rock 27.96 2.22
9(a) Host rock 27.93 2.00
9(b) Host rock 27.57 2.02
11 Vein (#2) 26.91 2.01
11 Vein (#4) 27.50 2.00
11(2) Vein (#2) 27.80 1.96
11(1) Vein (#2) 28.25 1.62
11(3) Vein (#2) 27.85 1.81
11(4) Vein (#2) 28.03 1.80
11 Host rock 27.75 2.34
11 Host rock 27.45 1.90
12 Vein (#2) 24.28 1.88
13 Vein (#4) 27.14 0.75
Southern slope
14 Vein (#2) 24.64 3.20
14 Host rock 25.05 3.30
15 Vein (#4) 25.32 3.08
15(2) Vein (#4) 25.33 2.00
Wadi Nakhar
16(2) Vein (#3) 27.63 2.68
17 Vein (#2) 27.78 2.67
18 Vein 25.87 3.62
18 Vein 27.24 2.66
18(8) Vein 28.00 2.63
18(5) Vein 27.92 2.78
18(7) Vein 27.43 2.43
18(3) Vein 27.77 2.66
18(6) Vein 27.68 2.73
18(4) Vein 27.52 2.55
18(2) Vein 28.12 2.49
18(1) Vein 27.79 2.62
18 Host rock 27.86 2.24
19 Vein (#2) 26.85 2.03
174 C. HILGERS et al.
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though precipitation from light rain events can have d18O
values up to +8& due to evaporative loss during rainfall
(Clark & Fritz 1997, p. 51). Groundwater in some uncon-
fined alluvial and carbonate aquifers of northern Oman
have d18O values similar to local precipitation and range
from )5 to +3&. The higher (more enriched) values are
due partly to surface and subsurface evaporative loss (Clark
& Fritz 1997, pp. 88–89). Soil gas CO2 from eight localit-
ies in northern Oman have d13C values between )22 and
+8, but the majority are between )18 and )14& (Clark
1987). Sixteen carbonate-cemented alluvium, travertine,
and aeolianite samples from northern Oman that were ana-
lyzed by Clark (1987) have d18O values between 24 and
33&, and d13C values between )12 and +1&. These car-
bonate values are in approximate equilibrium with local
precipitation, shallow groundwater, and soil gas CO2 at
surface temperatures.
The isotopic values of veins #1–#5 in Jabal Akhdar are
consistent with vein formation in (i) fluids isotopically buf-
fered by the carbonate-dominated stratigraphy or (ii) relat-
ively pristine (unaltered) meteoric water with low dissolved
CO2 concentrations at low temperatures (less than
approximately 50�C). Such low temperatures during for-
mation of these veins are not consistent with the inferred
Table 2 Continued.
19 Host rock 27.73 2.81
20 Vein (#6) 28.07 2.71
21(1) Vein (#6) 27.28 1.57
21(2) Vein (#6) 27.43 1.58
21(3) Vein (#6) 27.67 1.65
21 Vein (#6) 27.38 1.61
22 Vein (#6) 27.58 1.66
23 Vein (#2) 27.62 1.80
23 Host rock 27.45 1.46
24(a) Vein (#2) 26.50 1.72
25 Vein (#3) 27.45 1.95
26 Vein (#6) 26.13 2.28
27 Vein (#6) 20.83 1.59
28 Vein (#6) 19.40 1.81
29 Vein 26.17 2.68
29(2) Vein 21.96 1.36
29 Host rock 25.41 2.30
30(1) Vein (#2) 22.22 1.85
30(2) Vein (#2) 24.09 1.98
30 Vein (#2) 24.26 2.03
30 Host rock 25.81 2.22
31 – 18.89 1.72
31(1) Vein 26.40 1.33
31 Vein 26.47 0.89
33 Host rock 25.03 1.63
34 Host rock 25.49 1.06
35 Host rock 25.93 1.90
36 Vein (#2) 27.65 1.83
36 Vein (#2) 27.06 0.17
36(2) Vein (#2) 26.92 0.06
37 Vein (#6) 25.13 1.79
38 Vein 27.59 2.04
38 Host rock 28.03 1.94
39 Host rock 26.95 0.50
40 Vein (#6) 24.13 2.02
41 Host rock 27.12 1.74
42 Vein (#4) 27.35 1.38
43 Vein (#4) 27.52 1.44
43 Host rock 25.41 1.26
43 Host rock 27.25 1.55
44 Vein (#6) 18.50 1.85
45(3) Vein (#5) 25.11 2.08
45(5) Vein (#5) 25.37 1.85
45(6) Vein (#5) 25.05 1.92
45(4) Vein (#5) 25.24 1.91
45 Vein (#5) 24.54 2.32
45 Host rock 25.48 2.17
46 Vein (#4) 27.81 1.71
46 Host rock 27.58 1.67
47 Vein (#1) 28.04 1.22
47 Host rock 28.35 2.06
48 Vein (#4) 25.83 1.82
48 Vein 26.38 1.73
49(3) Vein (#2) 28.05 2.10
49(1) Vein (#2) 29.30 2.88
49(1) Vein (#2) 28.44 2.49
49 Vein (#2) 29.23 2.87
49 Host rock 28.34 2.20
50 Vein (#2) 26.56 2.18
50 Host rock 27.75 2.24
0
1
2
3
4
15 20 25 30
0
1
2
3
4 Wadi Nakhar
Southern slope
Wadi Bani Awf
Wadi Mistal
Host rocks
Veins
d
18Od 13
Cd 13
CFig. 5. d18O and d13C values of host rocks and veins. Isotopic values of
host rock samples are similar to those documented in other mid-Cretaceous
sediments of the paleo-Tethys. Isotopic values of many vein samples are
similar to host rocks, though some veins have much lower d18O values.
The d18O variations are inferred to be the product of variable fluid-rock
interactions and isotopic exchange prior to vein formation.
Fracture sealing and overpressures in limestone 175
� 2006 The Authors
Journal compilation � 2006 Blackwell Publishing Ltd, Geofluids, 6, 168–184
burial history of Jabal Akhdar. We propose that the fluids
responsible for formation of vein sets #1–#5 were either
highly evolved meteoric water or formation-like fluids that
were isotopically buffered by the carbonate host rocks and
trapped in the limestones during formation of generation
#2 and later veins (see discussion below). Some veins of #6
and #7 generations with d18O values below approximately
22& potentially precipitated from more pristine meteoric
waters at temperatures above ca. 80�C or from fluids that
had isotopically exchanged with silicates. Although no
source(s) of fluids can be uniquely identified, the structural
evolution of Jabal Akhdar during formation of vein sets #6
and #7 are consistent with uplift and exhumation, which
would have brought the dome in closer proximity to shal-
low groundwater aquifers. This would be consistent with a
change in fluid pressure from near lithostatic to hydrostatic
conditions between formation of vein sets #5 and #7 (see
discussion below).
The relatively high d13C values obtained in the veins of
this study are not consistent with significant amounts of
hydrocarbons having been involved in vein formation. This
does not preclude, however, the potential contribution
that minor hydrocarbon production might have had locally
on the formation of high fluid pressures.
DISCUSSION
A complex deformation history was responsible for the for-
mation of these seven vein generations. The first phase of
deformation caused vertical shortening and formation of
bedding-parallel stylolites with bedding-normal veins (#1).
This probably occurred during basin subsidence when the
maximum principal stress r1 was vertical and equal to the
overburden stress (r1¼qgz, where q is the density, g the
gravitational acceleration, z the thickness of the overbur-
den). Assuming the stylolites started forming at a depth of
1.5–2 km (cf. Bjørlykke 1989), the overburden would have
been approximately 50 MPa and effective vertical stress
would have been 30 MPa, if pore-fluid pressure was equal
y = 0.99x
R
2 = 0.78
22
24
26
28
30
22 24 26 28 30
y = 0.99x
R
2 = 0.66
1
2
3
4
1 2 3 4
d 13Cwall
d 13C
vein
d 18Owall
d 18O
vein
#1#2#3#4#5
#1#2#3#4#5
Fig. 6. d18O and d13C data of early-formed veins (sets #1–#5) and adja-
cent host rocks pairs are similar, consistent with a host-rock-buffered sys-
tem. The samples’ error bar of ±0.2& parallel to the y-axes represent the
largest intra-vein variation documented in this study (cf. Fig. 7). The stand-
ard deviations of our analytical measurements are 0.15& for d18O and
0.05& for d13C, and thus are similar in size to each datum symbol.
1.0
1.4
1.8
2.2
2.6
3.0
0 25 50 75 100
Across vein #11 (set #2)
Across vein #18 (set #2)
Across vein #21 (set #6)
Along vein #45 (set #5)
Along vein #49 (set #2)
24
26
28
30
0 25 50 75 100
d 13C
d 18O
Normalized distance
Fig. 7. Isotopic variations in individual veins documented normal (across-
vein) and parallel (along-vein) to the vein walls. The very limited variation
in isotopic values within the veins on a centimeter scale is consistent with
the fluid remaining isotopically constant during formation of individual
veins.
176 C. HILGERS et al.
� 2006 The Authors
Journal compilation � 2006 Blackwell Publishing Ltd, Geofluids, 6, 168–184
to hydrostatic pressure. The differential stress would have
been approximately 18 MPa if it is assumed that the effect-
ive horizontal stress, �0h ¼ �0
3, can be inferred using the
coefficient of earth pressure at rest, K0, to be 0.4 and
equal to �0h=�
0v ¼ K0, which takes into account the inelastic
behavior of rock due to compaction and diagenesis (Jones
et al. 1991; Mandl 2000, p. 179; Fig. 9A).
Formation of extensional veins #2 is consistent with r1
having been vertical, similar to its orientation during forma-
tion of veins #1 and basin subsidence. The limestones host-
ing veins #2 must have been cohesive as the fractures were
planar and transgranular. Formation of tensile fractures
must have required a higher internal fluid pressure
pf > rN + T, at a low differential stress r1)r3 £ 4T and
r3 ¼ rN ¼ )T (rN is the stress acting normal to the frac-
ture surface; T the tensile strength of the rock). Assuming
K0 to be 0.4 and a tensile strength of 10 MPa for limestone
(Thuro et al. 2001), the formation of the tensile fractures
would have required a differential stress of <40 MPa, which
would correspond to an overburden of about 4.5 km depth
and a fluid overpressure of 36 MPa (Fig. 9B). The fluid
pressure and the differential stress change can be calculated
if we assume a consolidated rock, which may be approxima-
ted as an elastic material. The horizontal stress is derived as
rh¼[m/(1)m)]rv. Assuming poroelastic deformation, the
effective horizontal stress decreases more slowly than the
pore pressure increases, causing a decrease in differential
stress with increasing overpressure (Engelder 1994; Engeld-
er & Fischer 1994; Mandl 2000, pp. 165–188) (Fig. 9C).
This can change the overall failure mode and promote ten-
sile over shear failure during burial of sedimentary basins at
greater depth (Hillis 2001). In isotropically and laterally
confined layers, the horizontal stress is increased by rh¼[m/(1)m)]rv + a[(1)2m)/(1)m)]pf (e.g. Segall 1989; Teufel
et al. 1991; Engelder 1994; Engelder & Fischer 1994;
Mandl 2000, pp. 165–188). The term m is the Poisson ratio
(0.25 dimensionless), and a is the Biot coefficient of effect-
ive stress (a ¼ 0.5 for strong rock, Mandl 2000, p. 171).
This would be consistent with the formation of tensile frac-
tures below 4.5 km.
We interpret the pinch-and-swell veins (#3) to have ini-
tially formed as tensile fractures that rotated and deformed
during subsequent shearing. If true, the nucleation of these
tensile fractures could only have occurred if the fluid was
overpressured. Some of these quartz-calcite veins were
filled and brecciated multiple times with different fluids.
Angular vein clasts floating in vein matrix are consistent
with these veins having formed as dilation breccias (Sibson
d 13C
d 18O
#1 & #2BE
#5 #6 #7#4#3
Host rock #1 stylolite veins#2 extension veins Eboudin veins B
#3 pinch and swell#4 parallel veins
#5 tension gashes #6 normal fault veins
#7 thrust related veins
15
20
25
30
0
1
2
3
4
Wadi NakharSouthern slopeWadi Bani AwfWadi Mistal
Fig. 8. Stable isotope values of host rocks and
seven sets of veins. The system became open to
external fluids during the latest vein sets, i.e.,
veins precipitated in dilational jogs along normal
faults (#6) and thrust faults (#7).
Fracture sealing and overpressures in limestone 177
� 2006 The Authors
Journal compilation � 2006 Blackwell Publishing Ltd, Geofluids, 6, 168–184
1985; Tarasewicz et al. 2005). Vein set #3 is not truncated
by cleavage surfaces, and thus potentially formed contem-
poraneously with the cleavage.
Four general models have been formulated to explain
the formation of bedding-parallel veins similar to vein set
#4. According to one model (i) (e.g. Henderson et al.
1990; Cosgrove 1993; Jessell et al. 1994), supra-hydrosta-
tic fluid pressure and a reorientation of r1 from bedding-
normal to bedding parallel could have formed tensile veins
#4.1 parallel to bedding. Alternatively, dilational exten-
sion-shear (or mixed-mode) fractures may have formed
oblique to r1, if the differential stress was between 4 and
5.66 T (e.g. Sibson 1998, 2000; Schultz 2000). In either
case, the fluid pressure would have been supra-hydrostatic.
According to a second model (ii) (e.g. Cox 1987; Koehn
& Passchier 2000), veins can form by the coalescence of
dilational jogs between adjacent shear surfaces. It has been
shown in some studies that such veins contain inclusions
40
–20 0 20 40 60 80 100 120s 'h s ‘ns 'v–T
K0 = 0.4
z = 2000 m
A
20
40
0 20 40 60 80 100 120–20 80 120s ‘h s 'v s ‘n–T
K0 =0.4
z = 4400 m
n = 0.25
B
20
40
–20 0 20 40 60 80 100 120
D
s 'h s 'v s ‘n
Fluid pressure increases duringreorientation of principal stresses
z = 4400 m
40
20
–20 0 20 40 60 80 100 120 140 160 180s 'h s 'hs 'v s ‘n
z = 4400 m
Fluid pressure remains constant duringreorientation of principal stressesE
20
40
–20 0 20 40 60 80 100 120 s ‘n
z = 8000 m
K0 = 0.4n = 0.25
s s s ss
s
s s
s s
s ss 'hs 'v
F
20
40
–20 0 20 40 60 80 100 120s 'h s 'v s ‘n
Poroelasticreduction of s1–s3
K0 = 0.4z = 4400 m
C
a = 0.5n = 0.25
Fig. 9. Mohr circle evolution to generate the observed vein pattern. (A) Stylolite veins (#1) formed early during basin subsidence by dissolution-precipitation
processes, when the sediments were unconsolidated or slightly consolidated. Increase in rock strength is indicated by changing color of the failure envelope
from gray to black. The gray Mohr circle displays the stress conditions at a hydrostatic fluid pressure at 2 km depth. The black Mohr circle shows the condi-
tions during tensile failure for vein set #1 formation. The arrow outlines the fluid overpressure required to form tensile fractures. (B) Extensional veins (#2)
also formed during basin subsidence. The strength of the rock increased because of consolidation. In order to form extensional veins, the fluid pressure was
supra-hydrostatic. The gray Mohr circles show the stress conditions at hydrostatic fluid pressure, the black Mohr circle the conditions at tensile failure. The
corresponding arrow outlines the fluid overpressure required to cause failure. We plotted two Mohr circles at hydrostatic conditions (gray) to display the dif-
ference in differential stress for elastic and inelastic stress conditions (using a Possion ratio of m ¼ 0.25 and K0 ¼ 0.4, respectively). The maximum differential
stress of 40 MPa is governed by the tensile strength of limestone. Note that the conditions at a depth of 4.4 km present tensile (K0 ¼ 0.4) and shear failure
(m ¼ 0.25), respectively. Once fractured, the fluid will drain and seal the fluid pathway, causing new overpressures. (C) An increase of fluid pressure may
cause a decrease in differential stress, if the deformation mechanism is poroelastic. This may cause tensile failure in settings where the total stresses generally
would cause shear failure. However, failure would require fluid overpressures in a porous rock being larger than expected at elastic conditions. A leakage of
the seal would release the overpressure, increase differential stress, and ultimately cause shear failure. (D) Bedding-parallel veins formed with the maximum
principal stress oriented at high angle to bedding. We consider two models: (i) The fluid pressure increased during re-orientation of the principal stress from
#2 to #4, (ii) the fluid pressure remained constant during reorientation of the principal stress. These are exemplified for hybrid extension shear (mixed mode)
failure. Model (i) describes the switch of the principal stresses and formation of bedding-parallel veins while fluid pressure increases. An increase in fluid pres-
sure caused a reduction of the effective vertical stress �0v until the differential stress became zero. (E) Model (ii) forces the differential stress to decrease until
the horizontal stress equals the vertical stress at constant fluid pressure. Then the differential stress increased until shear failure fractured the rock (e.g. black
Mohr circle). Fluid overpressures may have caused failure at lower differential stresses. Veins form in dilational jogs, which connect during progressive shear.
(F) Veins associated with normal faults and thrusts formed at high differential stresses during shear failure. The graph sketches the stress conditions at 8 km
depth for different stress conditions, and displays the amount of overpressure required to cause failure. Our data do not allow to determine the actual depth
during shear failure.
178 C. HILGERS et al.
� 2006 The Authors
Journal compilation � 2006 Blackwell Publishing Ltd, Geofluids, 6, 168–184
trails that are parallel to the steps/ramps (Jessell et al.
1994; Koehn & Passchier 2000). According to a third
model (iii), which is a variation of model (ii), extensional
fractures can form under compressive effective confining
pressures by the alignment of conjugate en echelon exten-
sion arrays that form at differential stresses above 5.66 T
(Teixell et al. 2000, and references therein). And according
to a fourth model (iv), which is a variation of model (i),
veins may form normal to r1 if the difference in tensile
strengths parallel and normal to bedding or foliation is lar-
ger than the differential stress (i.e. r3 + Th > r1 + Tv where
Th and Tv are the horizontal and vertical tensile strengths,
respectively), and if pore fluid pressure is pf > r1 + Tv (Gra-
tier 1987; Cosgrove 1995; example given in Jolly & Lone-
rgan 2002). The inferred roles of fluid pressures vary
significantly among these models.
Veins #4.1 probably did not form according to model
(iii) as the veins are tens of meters long, rarely cut across
the surrounding host rock, and do not contain host rock
‘bridges.’ If we assume the bedding-parallel veins formed
by supra-lithostatic fluid pressures with the maximum prin-
cipal stress r1 acting horizontally (mechanism i), then the
differential stress must have decreased below ca. 40 MPa
during formation of vein sets #2–#4. The transition from
vein set #2 to #4 is characterized by a reorientation of the
principal stress axes in models (i) to (iii). This could either
occur during an increase of fluid overpressure (Fig. 9D),
resulting in randomly oriented tensile fractures at high
fluid pressures and low differential stress (Cosgrove 1995).
Such fractures have not been observed to be associated
with vein set #4.1. Or vein sets #2–#4 evolved at constant
fluid pressure (Fig. 9E). Host rock inclusion bands aligned
sub-parallel to the vein wall and a folded bedding-parallel
vein #4.1 indicate vein formation during shearing (mech-
anism ii) (see also Table 1). Furthermore, bedding-parallel
veins (#4.2) and sheared pinch-and-swell veins (#3) are
consistent with shear during vein growth, so that the maxi-
mum principal stress was oriented oblique to bedding.
Both vein sets #3 and #4 are not truncated by cleavage
planes, suggesting that they formed during compression
oblique to bedding at elevated fluid pressures. Vein set #4
formed after reorientation of the principal stress axes at
increased differential stress. Shear indicators and host rock
inclusions within the vein indicate that they formed by the
coalescence of dilational jogs during shear failure at eleva-
ted fluid overpressure (Fig. 9E).
En echelon vein arrays (#5) developed during nucleation
of brittle normal faults (e.g. Mazzoli & di Bucci 2003). As
vein sets #5 and #6 probably formed during one continu-
ous process, we propose these veins formed according to
model (iii) with r1 oriented normal to bedding. The fluid
pressure is interpreted to have dropped to almost hydrosta-
tic, consistent with the influx of meteoric water during
brittle normal faulting (Fig. 9F).
Veins #7 formed in dilational jogs along the thrusts,
which implies a rotation of the maximum principal stress
relative to bedding from an extensional (#5, #6) to a com-
pressional (#7) setting (Fig. 9F).
Cause for overpressures
The supra-hydrostatic fluid pressures in the limestones dur-
ing formation of vein sets #2–#5 might have been due to
stress-related processes (disequilibrium compaction, tec-
tonic stress), an increase in fluid volume (e.g. gas genera-
tion, thermal expansion of fluids), fluid movement (e.g.
transfer of pore pressure from overpressured rocks), and
stress-insensitive diagenetic processes (e.g. calcite cementa-
tion) (e.g. Plumley 1980; Swarbrick & Osborne 1998;
Nordgard Bolas et al. 2004). These processes might have
been accelerated during emplacement of the Samail ophio-
lite and Hawasina nappe on the Jabal Akhdar during late
Santonian and early Maastrichtian, which added a mini-
mum of 8 km of material on top of the carbonates (Breton
et al. 2004). Based on our stable isotope data, the supra-
hydrostatic fluid pressures were probably not due to signifi-
cant production of hydrocarbons, nor to long-range move-
ment from higher- to lower-pressure fluid cells. Vein sets
#1–#5 have formed due to stress-related processes and
cementation processes.
Fluid pressure decreased during uplift and exhumation
of the dome when normal faults (#6) cut through the
dome and increased the bulk permeability of the lime-
stones. Meteoric water invaded the dome’s carbonates and
precipitated veins #6 and #7.
Relation to regional geology
Several deformation phases have been documented in the
Mesozoic platform of the Jabal Akhdar in the late Creta-
ceous during early extension and later subduction and
obduction of the overlying Hawasina and Semail ophiolite
nappes (e.g. le Metour et al. 1990; Masse et al. 1997). The
veins’ orientations, which are consistent with at least two
rotations of the maximum principal stress from being normal
to bedding (#1, #2, #5, #6) to oblique to bedding (esp. #3,
#4, #7), display part of this complex deformation history.
The stylolites and bedding-normal veins (#1 and #2)
formed during basin subsidence. They may be related to
early conjugate normal faults that trend WNW–ESE to
NW–SE and do not continue into the upper Cretaceous
Muti formation (Rabu 1987; Boote et al. 1990). Breton
et al. (2004) related this extension with the pulling down
of continental lithosphere during subduction of autochtho-
nous Infracambrian basement and Mesozoic cover during
Turonian-Santonian time.
The next phase of regional deformation displayed by
bedding-parallel veins (#4) is northward directed bedding-
Fracture sealing and overpressures in limestone 179
� 2006 The Authors
Journal compilation � 2006 Blackwell Publishing Ltd, Geofluids, 6, 168–184
parallel shear on both the northern and southern flank of
Jabal Akdhar. Crosscutting relations with the pinch-and-
swell veins (#3) have not been observed, but this vein set
#3 can be associated with cleavage formation during pro-
gressive shearing.
Breton et al. (2004) described some normal faults that
formed prior to regional bedding-parallel slip in Wadi Mi-
stal. They correlated this slip to the formation of a regional
cleavage across Jabal Akdhar (see also le Metour 1988;
Rabu 1987). Consequently, either (i) the normal faults
Vein set #2
ConsolidationContinues
A
B
C
s1
s1
s1
Vein set #1
Consolidationduring burial
s1
Vein set #4
Vein sets #7
s3
s1
Vein set #5 and 6
Dep
th (
km)
Stress or pressure
Hydrostatic
Lithostatic
Dep
th (
km)
Stress or pressureD
epth
(km
)
Stress or pressure
Dep
th (
km)
Stress or pressure
Dep
th (
km)
Stress or pressure
#2#1
NS
#7
NNESSW
#4
#5,6
Effective normal stress
She
ar s
tres
s
Fig. 10. Cartoon of vein formation and their possible relation to the large tectonic setting. The left column qualitatively illustrates the stress evolution in
Mohr space and the corresponding fluid pressure evolution in a stress-depth graph (central column). (A) Stylolites (#1) form early during basin subsidence,
followed by vertical extension veins (#2). (B) Bedding parallel veins (#4) contain shear indicators showing top-to-the NNE shear, such as small northward ver-
ging folds, which may represent a phase of exhumation. They are transected by two vein sets (tension gashes, dilation jogs) related to normal faulting
(#5, #6), (C) Late thrust planes contain veins in dilation jogs (# 7) and may be related to the doming of Jabal Akdhar.
180 C. HILGERS et al.
� 2006 The Authors
Journal compilation � 2006 Blackwell Publishing Ltd, Geofluids, 6, 168–184
associated with veins #5 and #6 formed after regional clea-
vage formation because they are not off-set by bedding-
parallel slip in Wadi Nakhar, (ii) the offset of normal faults
by bedding-parallel slip in Wadi Mistal is not related to the
regional top-to-the NNE shear, but to a later slip system
(e.g. flexural slip folding of the dome), (iii) different gener-
ations of normal faults formed prior and after regional clea-
vage formation, or (iv) bedding-parallel veins in Wadi
Nakhar are not related to bedding-parallel slip. The latter
two arguments are unlikely because all vein systems (apart
from #3) are present on both the northern and southern
flank of Jabal Akdhar and shear criteria consistent with
bedding-parallel slip are exposed in all wadis.
Vein sets #3 and #4 formed when r1 was oblique to
bedding and may have been coeval with regional cleavage
formation. This interpretation is consistent with the results
of other studies that associated a regional top-to-NNE
shear with cleavage formation in Jabal Akdhar during the
Campanian and Early Maastrichtian (le Metour et al.
1990; Breton et al. 2004). le Metour et al. (1990) noted
that subhorizontal cleavage increased in intensity toward
the northeast of Jabal Akdhar, which they related to sub-
duction. A second obduction-related mylonitic cleavage
then formed, which only affected the rocks within a few
tens of meters of the thrust plane between autochthonous
rocks and the Hawasina nappes (le Metour et al. 1990). le
Metour et al. (1990) correlated the first cleavage with
north- to northeast-directed bedding-parallel shearing.
Breton et al. (2004) related top-to-NNE shear with the
exhumation of the more buoyant subducted autochtho-
nous Infracambrian basement and Mesozoic cover in a
compressive regime which would be consistent with vein
set #4. This could be associated with top-to-northeast
shearing in the Saih Hatat region during exhumation (Gray
et al. 2004b).
All these deformation events (#3, #4) are related to the
late Cretaceous to Paleocene (Eoalpine) compression of
the northern Oman Mountains. The relative timing of
veins that formed after the cleavage is unclear because the
Tertiary cover is completely absent in the Jabal Akhdar
dome. Apatite fission track data of Jabal Akdhar and Saih
Hatat are consistent with Eoalpine orogenesis during the
late Maastrichtian-Paleocene (Poupeau et al. 1998). The
data preclude late Tertiary extension associated with uplift
and erosion (Mann et al. 1990), but are consistent with an
isothermal period from Eocene to Miocene (Poupeau et al.
1998). A second compression phase started in the early
Miocene (Burdigalian) and culminated during the late
Miocene-Pliocene, which initiated the uplift of the Oman
mountains and deposition of a molasse (Barzaman forma-
tion) (Carbon 1996; Poupeau et al. 1998). The Pliocene-
Quarternary is characterized by vertical uplift and normal
faulting (Hanna 1986; Patton & O’Connor 1988; Mann
et al. 1990; Carbon 1996; Poupeau et al. 1998).
Subhorizontal grooves on veins (#6) are consistent with
reactivation of normal faults as strike-slip faults, but no
clear sequence of deformation events relative to vein set #7
can be identified in our data set. The thrust-related veins
(#7) probably formed during Miocene-Pliocene doming of
Jabal Akhdar, eventually coeval with the reactivation of
normal faults.
The different vein sets show several changes in the stress
regime and may be associated with different regional defor-
mation phases. According to the regional geology outlined
above, the earliest vein sets #1 and #2 formed during sub-
sidence, while vein sets #3 and #4 may be related to sub-
duction/exhumation. This is followed by normal faulting
associated with vein sets #5 and #6, which were reactivated
to accommodate strike-slip motion. Vein #7 formed during
a late contractional phase of deformation (Fig. 10).
CONCLUSION
The Jabal Akhdar dome contains multiple calcite vein sets.
Stylolite veins (#1) and subvertical veins (#2) formed dur-
ing subsidence, the latter having formed at supra-hydrosta-
tic fluid pressures. Locally, pinch-and-swell veins (#3)
record a complex deformation history that involved both
calcite and quartz cementation that formed during shear-
ing. Fluids were overpressured, as indicated by the breccia-
tion of quartz and calcite vein fragments. Vein set #4 also
formed during shearing with r1 oriented oblique to bed-
ding at moderate fluid pressures, resulting in the formation
of bedding-parallel veins (#4.1, #4.2). The relative timing
of formation of veins #3 and #4 is unknown. The stable
isotope data of calcite are consistent with the early veins
having formed in a rock-buffered environment, and that
fracture-controlled fluid flow over long distances and
across different lithologic units was probably not import-
ant.
The following extension phase is displayed as en echelon
vein arrays (#5) and normal faults (#6). During normal
faulting, fluid pressure decreased and meteoric fluids infil-
trated the dome. The maximum principal stress r1 rotated
from being normal to sub-parallel to bedding, resulting in
the formation of thrusts and associated veins (#7) during
doming of Jabal Akhdar.
Further work will include a more systematic study of the
different vein generations and the heterogeneity of veins
associated with faults along-strike, and will better constrain
the pressure–temperature conditions during vein forma-
tion.
ACKNOWLEDGEMENTS
C. Hilgers would like to thank the German Academic
Exchange Service and the German Research Council (project
no. Hi-816/2) for funding parts of this project. Werner
Fracture sealing and overpressures in limestone 181
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Kraus is thanked for field assistance. Stimulating discussions
with Abdul Nassar Darkal (Sultan Qaboos University) and
Mohammed Alwardi (University of Leeds) are much appre-
ciated. We would like to acknowledge David Gray and Bob
Gregory for their reviews, which significantly improved the
manuscript.
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