fracture sealing and ﬂuid overpressures in …...fracture sealing and overpressures in limestone...

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: [email protected]. 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

� 2006 The Authors

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


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.



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




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.




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.




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




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.




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)



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.




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).




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.




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-




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.




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




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.

REFERENCES

Barnes HL, Rose AW (1998) Origins of hydrothermal ores. Sci-ence, 279, 2064–5.

Bechennec F, le Metour J, Rabu D, Villey M, Beurrier M (1988)The Hawasina basin: a fragment of a starved passive continental

margin, thrust over the Arabian platform during obduction of

the Sumail nappe. Tectonophysics, 151, 323–43.

Bechennec F, le Metour J, Platel JP (1993) Geological Map ofOman, scale 1:1,000,000. Ministry of Petroleum and Minerals,

Sultanate of Oman.

Beurrier M (1988) Geologie de la Nappe ophiolitique de Samail

dans les parties orientale et centrale des Montagnes d’Oman.PhD Thesis BRGM Orleans, 412 pp.

Beurrier M, Ohnenstetter M, Cabanis B, Lescuyer JL, Tegyey M,

Le Metour J (1989) Geochimie des filons doleritiques et des

roches volcaniques ophiolitiques de la nappe. Bulletin de laSociete Geologique de France, 8, 205–19.

Bjørlykke K (1989) Sandstone diagenesis and porosity modifica-

tion during basin evolution. Geologische Rundschau, 78, 243–68.

Bjørlykke K (1997) Lithological control on fluid flow in sedimen-

tary basins. In: Fluid Flow and Transport in Rocks. Mechanismsand Effects (eds Jamtveit B, Yardley BWD), pp. 15–34. Chap-man & Hall, London.

Boote DRD, Mou D, Waite RI (1990) Structural evolution of the

Suneinah foreland, Central Oman Mountains. In: The Geologyand Tectonics of the Oman Region (eds Robertson AHF, SearleMP, Ries AC), pp. 397–418. Geological Society, London.

Boudier F, Bouchez JL, Nicolas A, Cannat M, Ceuleneer G, Mis-

seri M, Montigny R (1985) Kinematic of oceanic thrusting inthe Oman Ophiolite: model for plate convergence. Earth andPlanetary Science Letters, 75, 215–22.

Boudier F, Ceuleneer G, Nicolas A (1988) Shear zones, thrusts

and related magamtism in the Oman ophiolite: Initiation ofthrtusting on an oceanic ridge. Tectonophysics, 151, 275–96.

Breton JP, Bechenne F, le Metour J, Moen-Maurel L, Razin P

(2004) Eoalpine (Cretaceous) evolution of the Oman Tethyan

continental marin: Insights from a structural field study in JabalAkdhar (Oman mountains). GeoArabia, 9, 1–18.

Burkhard M, Kerrich R (1988) Fluid regimes in the deformation of

the Helvetic nappes, Switzerland, as inferred from stable isotopedata. Contributions to Mineralogy and Petrology, 99, 416–29.

Carbon D (1996) Tectonique post-obudction des montagnes

d’Oman dans le cadre de la convergence Arabie-Iran. PhD The-

sis, Universite de Montpellier II, Montpellier, 286 pp.Clark ID (1987) Groundwater resources in the Sultanate of

Oman: Origin, circulation times, recharge processes and paleo-

climatology. Isotopic and Geochemical approaches. PhD Thesis,

Universite Paris XI, Paris, 264 pp.Clark ID, Fritz P (1997) Environmental Isotopes in Hydrogeology.

CRC Press, London, 328 pp.

Cosgrove JW (1993) The interplay between fluids, folds and

thrusts during the deformation of a sedimentary succession.Journal of Structural Geology, 15, 491–500.

Cosgrove JW (1995) The expression of hydraulic fracturing inrocks and sediments. In: Fractography: Fracture Topography as aTool in Fracture Mechanics and Stress Analysis (ed. Ameen MS),

pp. 187–96. Geological Society, London.

Cox SF (1987) Antitaxial crack-seal vein microstuctures and theirrelationship to displacement paths. Journal of Structural Geology,9, 779–87.

Cox SF, Knackstedt MA, Braun J (2001) Principles of structuralcontrol on permeability and fluid flow in hydrothermal systems.

Reviews in Economic Geology, 14, 1–24.

Engelder T (1994) Brittle crack propagation. In: ContinentalDeformation (ed. Hancock PL), pp. 43–52. Pergamon PressLtd, Oxford.

Engelder T, Fischer MP (1994) Influence of poroelastic behavior

on the magnitude of minimum horizontal stress, Sh, in over-

pressured parts of sedimentary basins. Geology, 22, 949–52.Ghisetti F, Kirschner DL, Vezzani L, Agosta F (2001) Stable iso-

tope evidence for contrasting paleofluid circulation in thrust

faults and normal faults of the central Apennines, Italy. Journalof Geophysical Research, 106, 8811–25.

Glennie KW (1995) The Geology of the Oman Mountains. Scientific

Press Ltd, Beaconsfield. 92 pp.

Glennie KW, Boeuf MGA, Hughes Clarke MW, Moody-Stuart M,Pilaar WFH, Reinhardt BM (1973) Late Cretaceous nappes in

Oman mountains and their geological evolution. American Asso-ciation of Petroleum Geologists, 75, 5–27.

Glennie KW, Hughes-Clarke MW, Boeuf MG, Pilaar WF, Rein-hardt BM (1974) Geology of the Oman mountains. Verhandel-ingen Koninklijk Nederlands geologisch mijnbouwkundigGenootschap, 31, 1–423.

Gratier JP (1987) Pressure solution-deposition and associated tec-

tonic differentiation in sedimentary rocks. In: Geological SocietySpecial Publication (eds Jones M, Preston RMF), pp. 25–38.

Geological Society, London.Gray DR, Gregory RT, Miller JM (2000) A new structural profile

along the Muscat-Ibra transect, Oman: Implications for

emplacement of the Samail ophiolite. In: Ophiolites and OceanicCrust: New Insights from Field Studies and the Ocean DrillingProgram (eds Dilek Y, Moores EM, Elthon D, Nicolas A), pp.

513–23. Geological Society of America Special Paper. Geologi-

cal Society of America, Boulder, CO.

Gray DR, Hand M, Mawby J, Armstrong RA, Miller JM, Greg-ory RT (2004a) Sm-Nd and Zircon U-Pb ages from the gar-

net bearing eclogites, NE Oman: Constraints on High-P

metamorphism. Earth and Planetary Science Letters, 222,407–22.

Gray DR, Miller JM, Foster DA, Gregory RT (2004b) Transition

from subduction- to exhumation-related fabrics in glaucophane-

bearing eclogites, Oman: evidence from relative fabric chronol-ogy and 40Ar/39Ar ages. Tectonophysics, 389, 35–64.

Gregory RT, Gray DR, Miller JM (1998) Tectonics of the Arabian

margin associated with the formation and exhumation of high-

pressure rocks, Sultanate of Oman. Tectonics, 17, 657–70.Guzzetta G (1984) Kinematics of stylolite formation and physics

of the pressure-solution process. Tectonophysics, 101, 383–94.

Hacker BR (1994) Rapid emplacement of young oceanic litho-sphere: argon geochronology of the Oman ophiolite. Science,265, 1536–65.

Hacker BR, Mosenfelder JL, Gnos E (1996) Rapid emplacement

of the Oman ophiolite: thermal and geochronological con-straints. Tectonics, 15, 1230–47.

Hanna SS (1986) The Alpine (Late Cretaceous and Tertiary) tec-

tonic evolution of the Oman mountains: a thrust tectonic




approach. In: Symposium on the Hydrocarbon Potential of IntenseThrust Zones, pp. 125–74. OAPEC, Kuwait.

Hanna SS (1990) The Alpine deformation of the Central Oman

mountains. In: Geological Society Special Publication (eds Ro-

bertson AHF, Searle MP, Ries AC), pp. 341–59. GeologicalSociety, London.

Henderson JR, Henderson MN, Wright TO (1990) Water-sill

hypothesis for the origin of certain veins in the Meguma Group,Nova Scotia, Canada. Geology, 18, 654–7.

Hillgartner H, van Buchem FSP, Gaumet F, Razin P, Pittet B,

Grotsch J, Droste H (2003) The Barremian-Aptian evolution of

the eastern Arabian carboante platform margin (northernOman). Journal of Sedimentary Research, 73, 756–73.

Hillis RR (2001) Coupled changes in pore pressure and stress in

oil fields and sedimentary basins. Petroleum Geoscience, 7, 419–

25.Hubbert MK, Rubey WW (1959) Role of fluid pressure in

mechanics of overthrust faulting. Bulletin of the Geological Soci-ety of America, 70, 115–66.

Jenkyns HC, Wilson PA (1999) Stratigraphy, paleoceanography,

and evolution of Cretaceous Pacific guyots: relics from a green-

house Earth. American Journal of Science, 299, 341–92.

Jessell MW, Willman CE, Gray DR (1994) Bedding parallel veinsand their relationship to folding. Journal of Structural Geology,16, 753–67.

Jolly RJH, Lonergan L (2002) Mechanism and controls on the

formation of sand intrusions. Journal of the Geological Society,159, 605–17.

Jones ME, Leddra MJ, Goldsmith AS, Yassir N (1991) Mecha-

nisms of compaction and flow in porous sedimentary rocks. In:Neotectonics and Resources (eds Cosgrove JW, Jones M), pp.

16–42. Belhaven Press, London.

Koehn D, Passchier CW (2000) Shear sense indicators in striped

bedding-veins. Journal of Structural Geology, 22, 1141–51.Konert G, Afifi AM, Al-Hajri SA, Droste H (2001) Paleozoic stra-

tigraphy and hydrocarbon habitat of the Arabian plate. GeoAra-bia, 6, 407–42.

Law BE, Ulmishek GF, Slavin VI (eds) (1998) Abnormal Pressurein Hydrocarbon Environments, Vol. 70. AAPG Memoir, Tulsa,

264 pp.

Mandl G (2000) Faulting in Brittle Rocks. Springer, London, 434

pp.Mann A, Hanna SS, Nolan SC (1990) The post-Campanian evolu-

tion of the Central Oman Mountains: Tertiary extension of the

Eastern Arabian margin. In: The Geology and Tectonics of theOman Region (eds Robertson AHF, Searle MP, Ries AC), pp.

549–63. Geological Society, London.

Masse J-P, Borgomano J, Al Maskiry S (1997) Stratigraphy and

tectonosedimentary evolution of a late Aptian-Albian carboantemargin: the northeastern Akdhar (Sultanate of Oman). Sedimen-tary Geology, 113, 269–80.

Masse J-P, Borgomano J, Al Maskiry S (1998) A platform-to-basin

transition for lower Aptian carbonates (Shuaiba Formation) ofthe northeastern Jebel Akhdar (Sultanate of Oman). SedimentaryGeology, 119, 297–309.

Mazzoli S, di Bucci D (2003) Critical displacement for normalfault nucleation from en-echelon vein arrays in limestones: a case

study from the southern Apennines (Italy). Journal of StructuralGeology, 25, 1011–20.

le Metour J (1988) Geologie del’autochtone des Montagnesd’Oman: la fenetre du Saih Hatat. PhD Thesis, Pierre and Marie

Curie University, Paris, 420 pp.

le Metour J, Rabu D, Tegyey M, Bechennec F, Beurrier M, Villey M

(1990) Subduction and obduction: two stages in the Eo-Alpine

tectonometamorphic evolution of the Oman Mountains. In: TheGeology and Tectonics of the Oman Region (eds Searle MP, Ries

AC), pp. 327–39. Geological Society, London.

Miller JM, Gray DR, Gregory RT (2002) Geometry and signifi-

cance of internal windows and regional isoclinal folds in north-east Saih Hatat, Sultanate of Oman. Journal of StructuralGeology, 24, 359–86.

Nicolas A, Ceuleneer G, Boudier F, Misseri M (1988) Structuralmapping in the Oman ophiolites: mantle diapirism along an

oceanic ridge. Tectonophysics, 151, 27–56.

Nordgard Bolas HM, Hermanrud C, Teige GMG (2004) Origin

of overpressures in shales: constraints form basin modeling.AAPG Bulletin, 88, 193–211.

Ortoleva PJ (ed) (1994) Basin compartments and seals. AAPGMemoir 61, 477 pp.

Patton TL, O’Connor SJ (1988) Cretaceous flexural history ofnorthern Oman mountain foredeep, United Arab Emirates.

AAPG Bulletin, 72, 797–809.

Plumley WJ (1980) Abnormally high fluid pressure: survey ofsome basic principles. AAPG Bulletin, 64, 414–30.

Poupeau G, Saddiai O, Michard A, Goffe B, Oberhansli R (1998)

Late thermal evolution of the Oman Mountains subophiolitic

windows: apatite fission-track thermochronology. Geology, 26,1139–42.

Pratt BR, Smewing JD (1993) Early Cretaceous platform-margin

configuration and evolution in the central Oman mountains,

Arabian peninsula. AAPG Bulletin, 77, 225–44.Putnis A, Mauthe G (2001) The effect of pore size on cementa-

tion in porous rocks. Geofluids, 1, 37–41.

Putnis A, Prieto M, Fernandez-Diaz L (1995) Fluid supersatura-tion and crystallization in porous media. Geological Magazine,132, 1–13.

Rabu D (1987) Geologie de l’autochtone des Montagnes d’Oman lafenetre du Jabal Akdhar . La semelle metamorphique de la Nappeophiolitique de Semail dans les parties orientale et centrale desMontagnes d’Oman: une revue. Universite Pierre et Marie Curie,

Paris 582 pp.

Ramsay JG (1980) The crack-seal mechanism of rock deformation.Nature, 284, 135–9.

Renard M (1986) Pelagic carbonate chemostratigraphy (Sr, Mg,

18O, 13C). Marine Micropaleontology, 10, 117–64.

Robb L (2005) Introduction to Ore-Forming Processes. Blackwell,Oxford.

Robertson AHF (1987) Upper Cretaceous Muti formation: trans-

ition of a Mesozoic nate plateform to a foreland basin in theOman mountains. Sedimentology, 34, 1123–42.

Rozanski K, Araguas L, Gonfiantini R (1993) Isotope patterns in

modern global precipitation. In: Climate Change in ContinentalIsotope Records (eds Swar PK, Lohmann KC, McKenzie J, SavinS), pp. 1–36. Geophysical Monograph. American Geophysical

Union, Washington D.C.

Scholle PA, Arthur MA (1980) Carbon isotope fluctuations in

Cretaceous pleagic limestones: potential stratigraphic and petro-leum exploration tool. AAPG Bulletin, 64, 67–87.

Schultz RA (2000) Growth of geologic fractures into large-

strain populations: review of nomenclature, subcritical crackgrowth, and some implications for rock engineering. Interna-tional Journal of Rock Mechanics and Mining Sciences, 37,

403–11.

Searle MP (1985) Sequence of thrusting and origin of culmina-tions in the northern and central Oman Mountains. Journal ofStructural Geology, 7, 129–43.

Searle MP, Warren CJ, Waters DJ, Parrish RR (2004) Structural

evolution, metamorphism and restoration of the Arabian




continental margin, Saih Hatat region, Oman Mountains. Jour-nal of Structural Geology, 26, 451–73.

Segall P (1989) Earthquakes triggered by fluid extraction. Geology,17, 942–6.

Sibson RH (1985) Stopping of earthquake ruptures at dilationalfault jogs. Nature, 316, 248–51.

Sibson RH (1998) Brittle failure mode plots for compressional

and extensional tectonic regimes. Journal of Structural Geology,20, 65–660.

Sibson RH (2000) Fluid involvement in normal faulting. Journalof Geodynamics, 29, 469–99.

Skelton PW (ed.) (2003) The Cretaceous World. Cambridge Uni-versity Press, Cambridge, 360 pp.

Skinner BJ (1997) Hydrothermal mineral deposits: what we do

and don’t know. In: Geochemistry of Hydrothermal Ore Deposits(ed Barnes HL), pp. 1–29. Wiley, New York.

Swarbrick RE, Osborne MJ (1998) Mechanisms that generate

abnormal pressures: an overview. In: Abnormal Pressures inHydrocarbon Environments (eds Law BE, Ulmishek GF, SlavinVI), pp. 13–33. AAPG Memoir, Tulsa.

Tarasewicz JPT, Woodcock NH, Dickson JAD (2005) Carbonate

dilation breccias: examples from the damage zone to the Dent

Fault, northwest England. Geological Society of America Bulletin,117, 736–45.

Taylor WL (1999) Fluid flow in discrete joint sets: field observa-

tions and numerical simulations. Journal of Geophysical Research,104, 28,983–9,006.

Teixell A, Durney DW, Arboleya M-L (2000) Stress and fuid con-trol on decollement within competent limestone. Journal ofStructural Geology, 22, 349–71.

Teufel LW, Rhett DW, Farrell HE (1991) Effect of reservoir

depletion and pore pressure drawdown on in situ stress anddeformation in the Ekofisk Field, North Sea. In: Rock Mechanicsas Multidisciplinary Science (ed. Roegiers JC), pp. 63–72. Balk-

ema, Rotterdam.Thuro K, Plinninger RJ, Zah S, Schutz S (2001) Scale effects in

rock strength properties. In: Part I: Unconfined Compressive Testand Brazilian test. Rock Mechanics – A Challenge for Society (eds

Sarkka P, Eloranta P), pp. 169–74. Balkema, Rotterdam, Espoo,Finland.

Tilton GR, Hopson CA, Wright JE (1981) Uranium-lead isotopic

ages of the Samail ophiolite, Oman, with application to Tethyan

ridge tectonics. Journal of Geophysical Research, 86, 2736–76.Warburton J, Burnhill TJ, Graham RH, Isaac KP (1990) The evolu-

tion of the Oman Mountains foreland Basin. In: The Geology andTectonics of the Oman Region (eds Robertson AHF, Searle MP,Ries AC), pp. 419–27. Geological Society, London.

Warren CJ, Parrish RR, Seare MP, Waters DJ (2003) Dating the

subduction of the Arabian continental margin beneath the Se-

mail Ophiolite, Oman. Geology, 31, 889–92.Weissert H, Lini A, Follmi KB, Kuhn O (1998) Correlation of

Early Cretaceous carbon isotope stratigraphy and platform

drowning events: a possible link? Palaeogeography, Palaeoclima-tology, Palaeoecology, 137, 189–203.




fracture sealing and ﬂuid overpressures in …...fracture sealing and overpressures in limestone...

Documents