Sedimentology and diagenetic evolution of the Neogene ‘Intermediate Sandstone Unit’ in the halite deposits of the Granada Basin (SE Spain): the turning point in the change from marine to continental sedimentation

Abstract

The Granada Basin is a small (50 × 50 km) Neogene intramontane basin located in the central part of the Betic Cordillera (Spain). In the latest Tortonian, the Granada Basin desiccated and a thick salt succession formed, encompassing three halite- bearing units: the ‘Lower Halite Unit’, the ‘Intermediate Sandstone Unit’ (ISU), and the ‘Upper Halite Unit’ (UHU). ISU deposits record the onset of marine to non-marine conditions in the Granada Basin. The main purpose of this paper is to study the environment of formation and the diagenetic evolution of the ISU salt-bearing unit, in order to assess the events leading to and resulting from the continentalization of this basin in the late Miocene. This study includes visual core descriptions, petrographic (conventional petrography and scanning electron microscopy), and geochemical (δ34S and δ18O and 87Sr/86Sr) analyses. ISU deposition took place in a coastal lake, isolated from the open sea by a sand barrier. Lake evolution was from a very shallow, hypersaline lacustrine environment to a deeper, perennial lake undergoing frequent storm-induced marine flooding and, finally, to a shallower, perennial saline lake. Isotope analyses point to a mixture of different inflow waters, including marine- and underground (hydrothermal)-water inputs for the origin of the brines. Halite dissolution occurred after flooding events and clear-halite cement was precipitated inside primary-halite dissolution cavities. Early diagenesis involves halite re-crystallisation during repetitive, dissolution–precipitation cycles, gypsum replacement by halite, halite replace- ment by nodular anhydrite, and framboid pyrite formation. Intermediate- to late-diagenetic processes are silica (megaquartz, chalcedony, and lutecite) replacement of halite and anhydrite, and celestine replacement. Megaquartz formation relates to sulfate-depleted, UHU percolating brines. Chalcedony and lutecite crystallization took place sometime later from sulfate- rich percolating brines, during deposition of the gypsum sequence occurring on top of the salt. Celestine, replacing lutecite, resulted from the interaction with Sr-rich underground waters (via dissolution of previously formed celestine).

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Journal of Iberian Geology (2018) 44:513–537
https://doi.org/10.1007/s41513-018-0069-z
RESEARCH PAPER
Sedimentology anddiagenetic evolution oftheNeogene
‘Intermediate Sandstone Unit’ inthehalite deposits oftheGranada
Basin (SE Spain): theturning point inthechange frommarine
tocontinental sedimentation
AdriánLópez-Quirós1 · MickaelBarbier2· JoséM.Martín3· XavierGuichet2
Received: 1 October 2017 / Accepted: 1 June 2018 / Published online: 12 June 2018
© Springer International Publishing AG, part of Springer Nature 2018
Abstract
The Granada Basin is a small (50 × 50km) Neogene intramontane basin located in the central part of the Betic Cordillera
(Spain). In the latest Tortonian, the Granada Basin desiccated and a thick salt succession formed, encompassing three halite-
bearing units: the ‘Lower Halite Unit’, the ‘Intermediate Sandstone Unit’ (ISU), and the ‘Upper Halite Unit’ (UHU). ISU
deposits record the onset of marine to non-marine conditions in the Granada Basin. The main purpose of this paper is to study
the environment of formation and the diagenetic evolution of the ISU salt-bearing unit, in order to assess the events leading
to and resulting from the continentalization of this basin in the late Miocene. This study includes visual core descriptions,
petrographic (conventional petrography and scanning electron microscopy), and geochemical (δ34S and δ18O and 87Sr/86Sr)
analyses. ISU deposition took place in a coastal lake, isolated from the open sea by a sand barrier. Lake evolution was from
a very shallow, hypersaline lacustrine environment to a deeper, perennial lake undergoing frequent storm-induced marine
flooding and, finally, to a shallower, perennial saline lake. Isotope analyses point to a mixture of different inflow waters,
including marine- and underground (hydrothermal)-water inputs for the origin of the brines. Halite dissolution occurred after
flooding events and clear-halite cement was precipitated inside primary-halite dissolution cavities. Early diagenesis involves
halite re-crystallisation during repetitive, dissolution–precipitation cycles, gypsum replacement by halite, halite replace-
ment by nodular anhydrite, and framboid pyrite formation. Intermediate- to late-diagenetic processes are silica (megaquartz,
chalcedony, and lutecite) replacement of halite and anhydrite, and celestine replacement. Megaquartz formation relates to
sulfate-depleted, UHU percolating brines. Chalcedony and lutecite crystallization took place sometime later from sulfate-
rich percolating brines, during deposition of the gypsum sequence occurring on top of the salt. Celestine, replacing lutecite,
resulted from the interaction with Sr-rich underground waters (via dissolution of previously formed celestine).
Keywords Granada Basin· Late Tortonian· Marine–continental transition· Evaporites· Diagenesis· Brine evolution
Resumen
La Cuenca de Granada (50 × 50km) es una cuenca neógena intramontañosa ubicada en la parte central de la Cordillera
Bética (España). Al final del Tortoniense se desecó, depositándose una potente secuencia salina, formada por tres unidades
que incluyen halita: la ‘Unidad de Halita Inferior’ (UHI), la ‘Unidad de Arenisca Intermedia’ (UAI) y la ‘Unidad de Halita
Superior’ (UHS). Los sedimentos de la UAI registran el cambio de condiciones marinas a no marinas. El objetivo principal
de este trabajo es estudiar el ambiente de formación y la evolución diagenética de la UAI, para evaluar los eventos que con-
dujeron y resultaron de la continentalización de esta cuenca en el Mioceno tardío. Este estudio incluye descripciones visuales
de los testigos de sedimento, análisis petrográficos (petrografía convencional y Microscopía Electrónica de Barrido) y geo-
químicos (δ34S, δ18O y 87Sr/86Sr). El depósito de la UAI tuvo lugar en un ambiente de lago costero, aislado del mar abierto
por una isla barrera. El lago evolucionó de somero e hipersalino, a más profundo y con frecuentes inundaciones marinas
inducidas por tormentas y, de nuevo a salino somero. Los análisis isotópicos indican mezcla de aguas, incluidas marinas
y subterráneas (hidrotermales) para explicar el origen de las salmueras. Los eventos de inundación disolvieron parte de la
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halita, reprecipitándola como cemento secundario en las cavidades de disolución. En la diagénesis temprana se produjo la
recristalización de halita, el reemplazamiento de yeso por halita y el de halita por anhidrita nodular, y la formación de pirita
framboidal. En la diagénesis intermedia-tardía tuvo lugar el reemplazamiento de halita y anhidrita por sílice (megacuarzo,
calcedonia y lutecita), y el de la lutecita por celestina. La formación de megacuarzo se relaciona con salmueras de percol-
ación empobrecidas en sulfato procedentes de la UHS. La cristalización de calcedonia y lutecita aconteció después a partir
de salmueras de percolación ricas en sulfato, procedentes de la secuencia yesífera que corona la sal. El reemplazamiento de
lutecita por celestina tuvo lugar por la interacción con aguas subterráneas ricas en Sr.
Palabras clave Cuenca de Granada· Tortoniense final· Transición marino-continental· Evaporitas· Diagénesis·
Evolución de la salmuera
1 Introduction
Evaporites as a salt rock are originally made up of minerals
that crystallized from brines saturated by solar evaporation
(Sarg 2001; Trichet etal. 2001; Warren 2006, 2016). Large
sequences of evaporites (saline crust and evaporite minerals)
form in arid/semiarid climatic settings and restricted basins
(Kendall and Harwood 1996; Schreiber and El Tabakh 2000;
War ren 2010). In recent years, evaporites have been studied
in relation to depositional environments (e.g. Taberner etal.
2000; Topper and Meijer 2013), mineralogy and geochem-
istry (e.g. Cendón etal. 2008; García-Veigas etal. 2013),
tectonic setting (e.g. Ortí etal. 2014a), and sequence stra-
tigraphy (e.g. Tucker 1991; Sarg 2001).
The study of evaporite units, based mainly on mineral-
ogical and/or geochemical studies, is usually complemented
by observations on the different sedimentological features
of the salt deposit (Schreiber and El Tabakh 2000). When
sediments still preserve their original mineralogy and mor-
phologies, the precise paleo-environmental conditions can
be inferred. Mineralogical studies in evaporite sequences
provide information on the type of basin (i.e. marine vs. con-
tinental, e.g. Strakhov 1970), and the principal salt sources
(solutes dissolved in seawater and recycling sources, e.g.
Lotze 1957). However, for a better understanding concerning
possible sources of original waters, geochemical analyses
are also needed. The isotope composition of sulfates leads
to the discerning of the marine vs. continental origin of the
SO2
−4 ion (i.e. marine vs. continental by recycling of older
evaporites, or by reoxidation of sulfides or elemental sul-
fur, e.g. Holser and Kaplan 1966; Nielsen 1972; Birnbaum
and Coleman 1979). Chemical parameters during evaporite
deposition as redox conditions can also be inferred from
sulfate isotopes (e.g. Pierre 1985, 1989). 87Sr/86Sr isotope
ratios have also been demonstrated to be a significant tool for
documenting the marine vs. continental origin of evaporites.
Studies linking mineralogical, petrological, sedimentologi-
cal and geochemical analysis provide a full understanding of
the depositional environment and paleohydrological evolu-
tion of evaporite basins (e.g. Lowenstein and Spencer 1990;
Ayo ra e tal. 1994a, 1995; Ortí etal. 2014a).
The Granada Basin, situated in the central part of the
Betic Cordillera (SE Spain), contains a thick, uppermost
Tortonian evaporite succession (the ‘Lower Evaporites’ of
Dabrio etal. 1982), formed under transitional marine to
continental conditions (Martín etal. 1984). This Evaporite
Unit, from the margin to the center of the basin, has the
following deposits: (a) gypsified stromatolites replaced by
celestine (the ‘Montevive and Escúzar Celestine’ sensu
Martín etal. 1984 and García-Veigas etal. 2015); (b) sel-
enite gypsum (the ‘Agrón Gypsum’ sensu García-Veigas
etal. 2015), and (c) halite (the ‘Chimeneas Halite’ sensu
García-Veigas etal. 2013) (Fig.1). In the salt succession
(up to 500m thick), García-Veigas etal. (2013) identi-
fied three halite-bearing units overlying a very thin, basal
anhydrite bed: the ‘Lower Halite Unit’ (LHU), the ‘Inter-
mediate Sandstone Unit’ (ISU) and the ‘Upper Halite Unit’
(UHU).
Fig. 1 Chemostratigraphic and lithostratigraphic correlation chart of
the Upper Miocene evaporites and celestine ore-deposits in the Gra-
nada Basin (modified from García-Veigas etal. 2015). 87Sr/86Sr and
δ34S isotope data are from García-Veigas etal. (2013, 2015). The
Lower Evaporites (Dabrio etal. 1982) comprise the Montevive and
Escúzar Celestine, the Agrón Gypsum, and the Chimeneas Halite
units (García-Veigas etal. 2015). The Alhama Gypsum (García-Vei-
gas etal. 2015) corresponds to the Upper Evaporites (Dabrio etal.
1982). A clastic unit (Cacín Lutite and La Malahá Turbidites; Dab-
rio etal. 1982) occurs in between. (1) Open-marine sediments; (2)
marine to continental, transitional deposits; (3–4) Lacustrine deposits
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Previous works have concentrated on the study of the
stromatolites and related celestine ore deposits (Martín etal.
1984; García-Veigas etal. 2015), on the gypsum (Dabrio
etal. 1982; Dabrio and Martín 1981), and on the halite
(García-Veigas etal. 2013), barely dealing with the interca-
lated ‘clastic’ sediments within the salt. This paper focuses
on the study of these latter deposits, linking petrological,
sedimentological, and isotope data to identify the environ-
ment of formation and the diagenetic evolution of this salt-
bearing unit. As shown below, the ISU represents the first
step in the growing influence of non-marine contributions to
salt deposition in the Granada Basin, so it is an appropriate
scenario for interpreting the events leading to and resulting
from the continentalization of this basin in the late Miocene.
2 Geological setting
The Betic Cordillera (southern Spain) is the westernmost
segment of the European Alpine Belt. Its paleogeographic
and tectonic evolution relates to the closure of the Tethys
Ocean in the Alpine Orogeny, during the Cenozoic (Sav-
ostin etal. 1986; Ziegler 1988; Dewey etal. 1989; Jabaloy
etal. 2002). Differential uplifting during the Alpine Orog-
eny produced a number of sedimentary basins within the
Betic Chain referred to as ‘Neogene Basins’. Some of these
basins opened directly to the Mediterranean Sea, while oth-
ers maintained links with the Atlantic Ocean through the
Guadalquivir foreland basin (Braga etal. 2003). The con-
nections between the Mediterranean-linked basins and the
Atlantic-linked basins were limited to a few seaways that
progressively closed in the course of the late Miocene (Mar-
tín etal. 2014).
Two different types of Mediterranean-linked basins can
be distinguished: the ‘inner basins’ (located far from the
present-day Mediterranean Sea) and the ‘outer basins’ (near
to the present-day Mediterranean Sea). The former group
includes, from West to East, the Granada, Guadix-Baza, For-
tuna and Lorca basins, while the latter group comprises the
Tabernas, Almería-Níjar, Sorbas and Alicante-Bajo Segura
basins (Braga etal. 2003) (Fig.2a). These basins were pro-
gressively isolated from the Mediterranean Sea (from late
Tortonian onward), until arriving at the present-day geo-
graphical situation (Sanz de Galdeano 1990; Sanz de Gal-
deano and Vera 1991, 1992; Braga etal. 2003; Sanz de Gal-
deano and Alfaro 2004; Krijgsman etal. 2006; Corbí etal.
2012). In the latest Tortonian–early Messinian, the inner
basins were disconnected from the Mediterranean Sea and
became continental, while the outer basins maintained their
links, in some cases up to the Pliocene (Braga etal. 2003).
The Granada Basin is a small (50×50km) Neogene intra-
montane basin located in the central part of the Betic Cordil-
lera (Fig.2). The basin’s sedimentary infill unconformably
overlies an irregular, fault-controlled, basement paleorelief
(Morales etal. 1990), consisting of rocks from the two major
domains of the Cordillera: the Internal Zones (cropping out
at Sierra Tejeda, Sierra de la Pera, and Sierra Nevada) and
the External Zones (cropping out at Sierra Gorda and Sierra
Arana) (Fig.2b). A series of sedimentary units can be differ-
entiated in the Neogene–Quaternary infilling of the Granada
basin (Martín etal. 1984; Braga etal. 1990, 2003) (Fig.2b).
Major fault systems have E–W orientations (Sanz de Gal-
deano 2008). Secondary faults, with a NW–SE trending, cut
and displace the E–W faults and define the principal subsid-
ing areas of the central and eastern part of the Granada Basin
(Rodríguez-Fernández and Sanz de Galdeano 2006).
The Granada Basin as such differentiated in the late Torto-
nian (at around 8.3Ma: Braga etal. 2003; Rodríguez-Fernán-
dez and Sanz de Galdeano 2006; Corbí etal. 2012). Initially,
the basin was a marine embayment connected to the Atlantic
Ocean to the northwest (Martín etal. 2014) and to the Mediter-
ranean Sea to the south and west (Braga etal. 1990). From 8.3
to 7.3Ma, major tectonic activity took place in the northeast-
ern (Sierra Arana) and eastern (Sierra Nevada) highland edges
of the basin (Fig.2), resulting in the deposition of significant
amounts of conglomerates at the base of the uplifted areas
(Braga etal. 1990, 2003; Martín and Braga 1997). Skeletal
carbonates accumulated in siliciclastic-free areas on platforms
around the marine-basin margins. Temperate-water carbonates
(Puga-Bernabéu etal. 2008; López-Quirós etal. 2016) formed
first, between 8.3 and 7.8Ma (Corbí etal. 2012), followed by
tropical, coral-reef carbonates (Braga etal. 1990), between
7.8 and 7.3Ma (Corbí etal. 2012). In the course of the late
Tortonian, the marine connections to the Atlantic Ocean were
interrupted first (Martín etal. 2014) and, finally, those to the
Mediterranean Sea (Martín etal. 1984) due to a major regres-
sion resulting from a significant eustatic sea-level fall associ-
ated with local tectonic uplift. As a result the Granada Basin
desiccated (Martín etal. 1984) and became continental in the
latest Tortonian (7.3–7.2Ma, Corbí etal. 2012).
During the Messinian and the Pliocene, the continental Gra-
nada Basin was filled by detrital (alluvial-fan and fluvial) and
carbonate/evaporite (lacustrine) deposits (Dabrio etal. 1982;
Martín etal. 1984; Fernández etal. 1996; García-Alix etal.
2008; García-Veigas etal. 2015; Fig.2b). During the Qua-
ternary, sedimentation concentrated in small, fault-controlled,
high-subsidence depocenters (Morales etal. 1990; Rodríguez-
Fernández and Sanz de Galdeano 2006; García-Alix etal.
2008), filled by detrital sediments (mostly conglomerates)
(Fig.2b).
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Fig. 2 Geological setting of the Granada Basin. (a) Major Neogene
sedimentary basins of the Betic Cordillera, Spain. (1): Alicante-
Bajo Segura Basin, (2): Fortuna Basin, (3): Lorca Basin, (4): Sorbas
Basin, (5): Tabernas Basin, (6): Almería-Níjar Basin, (7): Guadix-
Baza Basin, (8): Granada Basin, (9): Ronda Basin, (10): Guadalquivir
Basin (from Braga etal. 2003). (b) Simplified geological map and
Miocene to Quaternary stratigraphy of the Granada Basin (modified
from Dabrio etal. 1982; Martín etal. 1984). Color code for the map
is on the left side of the stratigraphic column. 1: Basement rocks; 2:
Miocene marine deposits; 3: Miocene marine to continental, transi-
tional (evaporite) deposits; 4: Miocene continental (lacustrine) depos-
its; 5: Pliocene–Quaternary continental (alluvial/fluvial) deposits.
P–M Palaeozoic–Mesozoic; Aq–lT Aquitanian–lower Tortonian; Pl–
Q Pliocene–Quaternary. Blue dot marks the position of the CMN-3
borehole
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3 Materials andmethods
3.1 Petrographic observations andfacies analysis
This study concentrates on the ISU deposits drilled in the
CMN-3 exploratory borehole, located close to the village of
Ven ta s d e H uelma (N 37° 06 ′46.3″, W3°47′20.3″; Figs.2, 3),
in the Granada Basin. A detailed log of the ISU is shown
in Fig.4. Twenty-seven rock samples were taken from a
55-m-thick interval (at depths of 545–600m).
Rock-facies analyses were based on visual (macroscopic)
and microscopic (thin section) observations. Sixteen thin
sections, about 30µm thick, were prepared and impreg-
nated by a blue epoxy (EpoBlue®) to distinguish pores from
textural components. Thin sections were examined under
polarized and cross-polarized light with a Nikon Eclipse LV
100 POL optical microscope. During the petrographic study,
several image sequences were taken with a camera (ProgRes
C10) connected to the microscope and captured with the
ProgRes 2.1 image-management program.
Fig. 3 Study area. a Simplified geological map of the Granada Basin
and b detailed 3D topographic mosaic showing the location of the
Lower Evaporites (sensu Dabrio etal. 1982) depositional area (1),
and of the subsurface halite (2) (Chimeneas Halite, sensu García-
Veigas etal. 2013, 2015). Note the position of the CMN-3 borehole
(yellow star). Nearby villages are shown on the map. c Simplified
geological cross-section of the Granada Basin and position of the salt
depocenter (modified from Rosino 2008)
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3.2 Mineralogical andgeochemical analyses
X-ray diffraction (XRD) was used to identify minerals as
well as to determine the average bulk composition. For XRD
determinations, 32 powders from ISU rocks were prepared
with an agate mortar, and then analyzed with an X’pert PRO
PW 3040/60 diffractometer operated at 50kV, 30mA. Sam-
ples were scanned in the 2θ range from 2° to 79°, with a
step size of 0.033° (2θ)/s. The counting time was 120min
per sample.
SEM examinations were performed on two carbon-coated
samples analyzed with a Zeiss Evo Ma 10 (10–15kV beam
current). Observations were carried out on backscattered
electron image (BEI) mode, used to examine minerals that
have initially been viewed and studied in conventional
petrography, in order to gain information not available from
the optical microscope.
Thirteen powders from discrete evaporite layers within
the ISU were sampled with a dental drill (Dremel 225), in
order to measure oxygen and sulfur stable-isotope composi-
tion (δ18OSMOW and δ34SCDT). Contamination between two
subsequently drilled samples was avoided by using diluted
HCl to dissolve the carbonate components from the drill
before each sampling. The isotope analysis was performed at
the CCiT-UB (Centres Científics i Tecnològics, Universitat
de Barcelona, Spain). To determine the isotope composi-
tion of the sulfate minerals (anhydrite), the samples were
dissolved in > 18MΩ cm−1 water, boiled, and recovered as
BaSO4 from a solution at pH ~ 3. The δ34SCDT was deter-
mined with a Carlo Erba 1108 elemental analyzer and the
δ18OSMOW with a TC-EA unit, both coupled to an IRMS
Thermo Finnigan Delta Plus XP. The analytical error (2σ)
was ± 0.2‰ for δ34S and ± 0.4‰ for δ18O (values for stand-
ard NBS-127: δ34S NBS-127: 20.3 ± 0.1‰; δ18O NBS-127:
9.3 ± 0.2‰).
Four evaporite (halite and anhydrite) powders were sam-
pled with a dental drill (Dremel 225), in order to measure
strontium isotopic ratios (87Sr/86Sr). The same procedure as
for the sampling of sulfate isotopes was followed to avoid
contamination between subsequently drilled samples. The
isotope analysis was performed at the SUERC (Scottish Uni-
versities Environmental Research Centre, Scotland, UK).
The measurements were carried out on 1mg of powder in
a 2.5M HCl. The separation between the two components
was done by the standard procedure of ionic exchange.
Samples were loaded onto single Re filaments and run on a
VG Sector 54e thermal ionization mass spectrometer. The
Sr ratio values were then normalized to a ratio value of
87Sr/86Sr = 0.1194. The precision was better than ± 0.04‰.
4 Results
4.1 Facies description
The term ‘facies’ is broadly used to refer to a set of charac-
teristics such as dimensions, sedimentary structures, grain
size and type, color and biogenic content of a sedimentary-
rock unit (Middleton 1973). The term ‘evaporite facies’ is
used here in the same sense as in Schreiber and Kinsman
(1975), Schreiber etal. (1976) or Aref etal. (1997) to refer
to a evaporite sediment/sedimentary rock deposited under
specific environmental conditions, regardless the age or
physiographic–stratigraphic position.
Six ISU facies, organized in vertically stacked, cm- to dm-
scale sequences throughout the study section (Fig.4), have
been recognized. They are classified into four groups: (1)
clastic lithofacies; (2) carbonate lithofacies; (3) mixed, car-
bonate–evaporite lithofacies; and (4) evaporite lithofacies.
4.1.1 Clastic lithofacies
(a) Bioclastic sandstone (C1 facies): This facies consists of
light to medium gray, moderately to well-sorted, fine- to
coarse-grained bioclastic sandstones exhibiting an overall
fining-upward trend (Fig.5a). The bioclastic sandstone lay-
ers are up to 20cm thick and intercalate between silts, car-
bonates, and evaporites (Fig.4). Basal contacts are sharp
and very irregular (Fig.5b). This facies locally exhibits
small (a few cm high) wave ripples (Fig.5c) and wave-rip-
ple cross lamination. Grains, ranging in size from 250μm
to 1cm, are of carbonate bioclasts, plant remains, angular
and highly irregular extraclasts (metamorphic and carbonate
rocks), K-feldspar, quartz, and reworked evaporitic grains
(Fig.5d–i). Bioclasts, usually fragmented, are from bryozo-
ans, bivalves, coralline algae, and foraminifera. Whole speci-
mens of globigerinidae (globigerines and globigerinoids),
globoratalia and other planktonic foraminifera (probably
orbuline) are frequently found. Some whole specimens of
benthic foraminifera (e.g. miliolids and elphidium) were also
observed. Halite often occurs as intra- and inter-granular
pore-filling cement (Fig.5e, i).
(b) Bioclastic siltstone (C2 facies): This massive to banded,
white to pale greenish-yellow silty facies occurs as tabular
beds 0.5–15cm thick, intercalated with bioclastic sands and
carbonates (Figs.4, 6a, b). Matrix consists of dense to pel-
loidal clay. Silt-sized grains are mainly from broken calcite,
feldspar and quartz crystals, and minor dolomite and glauco-
nite (Fig.6c). Highly fragmented, fossil remains (mostly from
planktonic foraminifera and bivalves) are also found (Fig.6c,
Fig. 4 a Condensed lithological log of the CMN-3 borehole (after
García-Veigas etal. 2013). b Detailed lithologic log of the Intermedi-
ate Sandstone Unit (ISU; CMN-3 borehole). Main lithologies, facies
and sedimentary structures are shown. Blue dots indicate the position
of the samples studied
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d). Plant debris (up to 0.6cm in width and 2cm in length) also
occur (Fig.6e, f). Halite is displayed as pore-filling cement.
4.1.2 Carbonate lithofacies: carbonate mudstone (CB1
facies)
This facies consists of light greenish-gray to dark brown-
ish-green massive carbonate mudstone that intercalates
with clastic and evaporite sediments (Figs.4, 7a). Car-
bonate beds range in thickness from a few mm up to
10cm. Calcite is the carbonate mineral, as determined by
X-ray diffraction. Small anhydrite nodules, up to 1cm in
size, are locally found. Millimeter-sized, displacive halite
crystals may also occur (Fig.7a). Terrigenous compo-
nents (calcite, feldspar and quartz grains, and clay miner-
als), if present, represent less than 15% of the rock bulk
Fig. 5 Bioclastic sandstone (C1 facies). a Core slab of a moderately
to well-sorted, fine- to coarse-grained sandstone exhibiting an over-
all fining-upward trend. Plant remains (yellow arrow) concentrate on
top. b Core slab of a thin, bioclastic sandstone layer (C1), intercalated
between carbonate mudstones (CB1), exhibiting a highly irregular
erosional base (yellow arrow). c Plan view showing poorly preserved
wave ripples (aligned roughly parallel to the b axis of the picture) on
top of a bioclastic sandstone layer. d–i Plain-polarized light (PPL)
photomicrographs of the bioclastic sandstones. Grains, up to 1cm
in size, are of carbonate bioclasts (planktonic and benthic foraminif-
era, bryozoans, bivalves, and coralline algae), plant remains, quartz,
K-feldspar, extraclasts (carbonate and metapelitic metamorphic
rocks) and reworked evaporite minerals. Foraminifer shells are seen
in d and e. A bivalve shell (yellow arrow), exhibiting a homogene-
ous prismatic wall structure, is shown in f. A coralline algal fragment
(yellow arrow), exhibiting a series of light and dark bands and a cel-
lular structure parallel to the long axis of the grain, can be found in g.
Extraclasts are observed in h. They are metamorphic dolomite (yel-
low arrow) and metapelitic clasts from the Alpujárride Complex. e
and i display intra- (yellow arrows) and inter-granular (white arrows)
pore-filling halite cement. Remaining porosity is shown in blue
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volume. Fossil remains, such as planktonic foraminifera
and plant debris (Fig.7b), and minor brachiopod/bivalve
shell fragments, are locally found.
4.1.3 Mixed, carbonate–evaporite lithofacies: carbonate/
anhydrite interbeds (CE1 facies)
This facies consists of an irregular alternation between
light brownish-yellow carbonate mudstone bands, up to
3cm thick, and white (cloudy) anhydrite laminae, rang-
ing in thickness from < 1mm to a few mm. The carbon-
ate/anhydrite ratio is highly variable, although carbonate
bands predominate. Lamination is commonly displayed as
plane-parallel, tabular structures (Fig.7c), but can also be
wavy in shape (Fig.7d). Anhydrite laminae show poorly
delineated, ghost relics of presumably former micro-sel-
enite gypsum crystals (Figs.7e, f) and are made up of
prismatic, disoriented lath crystals with a felted texture
(e.g. Maiklem etal. 1969; Holliday 1973) (Fig.7g, h).
4.1.4 Evaporite lithofacies
(a) Nodular anhydrite (E1 facies): This facies is composed
of layers, up to 8cm thick, made up of a mosaic of tightly
packed, white anhydrite nodules, separated from one another
by a light-brownish, carbonate mud matrix (Fig.8a). Single
nodules, ranging from a few mm to several cm in size, may
exhibit a subtle, chicken-wire internal structure (e.g. Warren
and Kendall 1985; Warren 2006, 2016). They consist of lath-
shaped anhydrite crystal aggregates with a felted texture.
(b) Banded halite (E2 facies): This facies consists of
small patches to centimeter- to decimeter-thick bands of
intercalated white (cloudy) halite and gray halite (Fig.8b,
c). Contacts with enclosing clastic and carbonate facies
(Fig.4) are usually sharp. The fine to coarse, up to 1cm in
size, cloudy halite crystals exhibit a chevron structure and
contain abundant fluid inclusions (Fig.8c–e), whereas the
coarse-grained, gray halite crystals are free of inclusions.
The latter display an equigranular, mosaic-type texture
(Fig.8f). Carbonate mud commonly occurs interstitially
Fig. 6 Bioclastic siltstone (C2 facies). a Core slab showing an ero-
sive, sharp contact between C1 and C2 facies (yellow arrow). b PPL
photomicrograph showing the contact between C2 and CB1 facies.
c PPL photomicrographs showing main textural components of C2
facies. Silt-sized grains are mainly from calcite, feldspar, quartz, car-
bonate bioclasts, and minor dolomite and glauconite (blue arrow).
Carbonate bioclasts include fragments of coralline algae, bivalves,
bryozoans (orange arrow) and foraminifera (red arrows). d PPL pho-
tomicrograph showing a bored, bivalve-shell fragment (white arrow).
Borings (yellow arrow) are filled by carbonate mud (micrite). e Rock
core sample (transversal section; view from above) of centimeter-
sized plant debris (yellow arrow). f PPL photomicrograph showing
pyrite particles (in black colour; yellow arrow) associated with plant
debris
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filling in the space between cloudy halite crystals as well
as filling in some fluid-inclusion voids (Fig.8e, g). All
this provides a chaotic, mottled appearance (also termed
haloturbated texture, e.g. Smith 1971; Holdoway 1978) to
the salt rock. Anhydrite nodules (up to 1cm in diameter)
concentrate at the base and the top of halite beds (Fig.8f,
h), often displayed as contorted bands of coalesced nod-
ules (enterolithic-like structures; e.g. Shearman and Fuller
1969; Warren and Kendall 1985; Warren 2006, 2016).
Anhydrite nodules are made up of lath-shaped crystals
with a felted texture.
4.2 Facies distribution
With respect to lithofacies distribution, a series of inter-
vals can be clearly distinguished in the ISU sequence
from bottom to top. The first interval (from 601 to 592m
deep; Fig.4) is characterized by the predominance of
Fig. 7 Carbonate (CB1) and mixed, carbonate–evaporite (CE1)
facies. a Core slab showing the contact between carbonate mudstone
(CB1) and banded halite (E2) facies. Millimeter-sized, displacive
halite crystals (blue arrow) are visible within CB1 facies. Millimeter-
sized anhydrite nodules (a), displaying an enterolithic-like structure,
occur inside the halite bed (yellow arrow). b Core slab showing a car-
bonate mudstone layer (CB1) containing abundant plant debris (yel-
low arrow). c, d Core slabs showing carbonate-dominated lithofacies
intercalating anhydrite laminae (yellow arrows) (CE1 facies). Anhy-
drite laminae are often displayed as plane-parallel, tabular structures
(as observed in picture c), but can also be wavy in shape (as reflected
in picture d). e PPL photomicrograph showing ghost relics of pre-
sumably former micro-selenite gypsum crystals (yellow arrow) within
a thin anhydrite lamina. f Cross-polarized light (CPL) photomicro-
graph showing in detail anhydrite-crystal pseudomorphs after micro-
selenitic gypsum (yellow arrows) in CE 1 facies. Anhydrite laminae
display prismatic, disoriented lath crystals with a felted texture (blue
arrow). g, h PPL and corresponding CPL photomicrographs showing
anhydrite laminae (from CE1 facies) made up of prismatic, disori-
ented lath crystals with a felted texture
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nodular anhydrite layers (E1 facies) and carbonate mud
(CB1 facies). In the second interval (from 592 to 564m
deep; Fig.4) clastic sediments (C1 and C2 facies), locally
intercalating halite (E2 facies) and carbonate/anhydrite
interbeds (CE1 facies), are dominant. In the third interval
(from 564 to 545m deep; Fig.4) evaporites, halite (E2
facies), and carbonate/anhydrite interbeds (CE1 facies),
are the most abundant deposits.
Fig. 8 Nodular anhydrite (E1) and banded halite (E2) facies. a Core
slab showing a mosaic of tightly packed, white anhydrite nodules
separated from one another by a carbonate matrix (E1 facies; yel-
low arrow). b Core slab displaying irregular bands of white (cloudy)
primary halite and gray mosaic-type, secondary halite (E2 facies).
Millimeter-sized anhydrite nodules (yellow arrow) occur within the
mosaic-type halite. c Core slab showing relics of primary, cloudy
halite crystals (yellow arrow), exhibiting a chevron structure, inside
patches of secondary, mosaic-type halite. d PPL photomicrograph
showing biphasic, fluid inclusions (yellow arrow) inside a primary,
white (cloudy) halite crystal. e PPL photomicrograph showing a close
view of the rows or clusters of fluid inclusion voids (yellow arrows),
filled by carbonate mud. Carbonate mud also fills the interstitial space
between halite crystals (blue arrows). f High-resolution image of a
scanned thin section showing upper contact of E2 facies. Mosaic-type
halite (H), millimeter-sized anhydrite nodules (yellow arrows), inter-
stitial carbonate mud between halite grains (M), as well as remains
of chevron structures (as rows or clusters of fluid inclusion voids; red
arrow), are indicated. g PPL photomicrograph showing equigranu-
lar, mosaic-type replacement halite. Note curved crystal boundaries
meeting at triple junctions (blue arrow). Rows or clusters of fluid-
inclusion voids, filled by carbonate mud (yellow arrow), delineate the
shape of the former, primary-halite crystal. h Core slab showing coa-
lesced anhydrite nodules, display as contorted bands (enterolithic-like
structures; yellow arrow), accumulated at the top of a halite band
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4.3 Additional petrographic observations: other
mineral occurrences
4.3.1 Gypsum (CaSO4·2H2O)
Gypsum occurs as single, lenticular to prismatic crystals,
up to 150μm in width and 400μm in length (Fig.9a, b),
preserved as halite pseudomorphs. It is common at the
base of some halite beds.
4.3.2 Pyrite (FeS2)
Pyrite occurs as framboid crystals, up to 100μm in size,
within the carbonate and clastic lithofacies (Fig.9c), com-
monly associated with plant debris (Fig.6f).
Fig. 9 Other mineral occurrences. a, b PPL and corresponding CPL
photomicrographs showing lenticular to prismatic gypsum crystals,
now preserved as halite pseudomorphs (yellow arrow), embedded in
carbonate mud (M). c Reflected light microscope (RLM) photomi-
crograph showing micron-sized pyrite framboids (in yellow colour;
white arrow) within carbonate mud. d, e PPL and corresponding CPL
photomicrographs showing partial replacement of an anhydrite nod-
ule (yellow arrow) by equigranular quartz (megaquartz; white arrow).
Relicts of the felted anhydrite texture can still be recognized in the
quartz. f CPL photomicrograph showing replacement of pre-existing
evaporites (i.e. halite and anhydrite) by zebraic chalcedony (yellow
arrow). g, h PPL and corresponding CPL photomicrographs show-
ing selected replacement of mosaic-type halite by lutecite at grain
boundaries (yellow arrow). i SEM photomicrograph (BSD) showing
replacement of mosaic-type halite (H) by lutecite (L), which in turn
was replaced by celestine (C; yellow arrow). Gypsum crystals, now
preserved as halite pseudomorphs (G; red arrow), are also visible
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4.3.3 Authigenic silica (SiO2)
Authigenic silica occurs as equigranular quartz (megaquartz)
and fibrous quartz (zebraic chalcedony and lutecite) replac-
ing evaporites, mainly halite and anhydrite. Equigranular
quartz consists of euhedral, equant crystals (> 20μm; mega-
quartz) within which felted anhydrite relicts are preserved
(Fig.9d, e). Chalcedony consists of sheaf-like bundles of
radial, length-fast fibers of quartz, in which the crystal-
lographic c-axis lies normal to fiber long axis. Low bire-
fringence yields first-order white/gray/yellow interference
colors. Spherulitic growth with maltese-cross or zebraic
extinction is observed in chalcedony (Fig.9f). Lutecite
consists of length-slow fibers of quartz (i.e. chalcedony),
in which c-axis is oriented at approximately a 30° angle to
the long axis of the fiber. High birefringence, second-order
yellow/magenta/blue interference colors were observed in
lutecite (Fig.9g–i).
4.3.4 Celestine (SrSO4)
Celestine occurs in very minor quantities as individual,
μm- to mm-sized crystals, associated with halite. This min-
eral was identified under scanning electron microscope
(SEM; Fig.9i), presenting a white color and lenticular
morphologies.
4.4 Isotope data
4.4.1 Sulfate isotopes (δ34S andδ18O)
Sulfur and oxygen isotope values from the ISU at the CMN-3
borehole are listed in Table1 and plotted as δ18OV-SMOW
vs. δ34SV-CDT in Fig.10. For comparison, values are also
shown from the Granada Basin (Chimeneas Halite Succes-
sion; García-Veigas etal. 2013) and other Iberian Peninsula
Miocene basins (Calatayud, Tajo, and Ebro basins; Utrilla
etal. 1992), along with values for Miocene marine sulfates
(from Claypool etal. 1980; Paytan etal. 1998), upper Tri-
assic gypsum deposits (Keuper; García-Veigas etal. 2013)
and Tortonian gypsum deposits (Rouchy and Pierre 1979).
The δ34SV-CDT values from the ISU range from +10
to +20.9‰ with an average of +18.5 ± 1.9‰ (Fig.11).
The lowest value (+10‰) was found in a nodular-anhy-
drite sample rich in pyrite. All these values are lower than
those estimated for Miocene marine sulfates (δ34S: +21 to
+24‰; Paytan etal. 1998) and higher than those recorded
from upper Triassic gypsum outcropping in the Granada
Basin (+14.3 ± 0.7‰; García-Veigas etal. 2013), with the
exception found at 556.8m deep (nodular-anhydrite sample
rich in pyrite). Some of these values are also comparable to
those reported for the Tortonian gypsum from the Granada
Basin (+16.7 ± 0.7‰; Rouchy and Pierre 1979).
Table 1 Geochemistry-analysis results of ISU samples: sulfate iso-
tope composition (δ18OV-SMOW; δ34SV-CDT), and 87Sr/86Sr isotope
ratios
Isotope values reported by García-Veigas etal. (2013) for the basal
anhydrite bed, and the LHU, ISU, and UHU, respectively, are also
shown for comparison
Depth (m) δ18O (‰) δ34S (‰) 87Sr/86Sr Source
Upper Halite Unit—UHU
382 16.2 17.2 García-Veigas etal.
(2013)
391.4 0.708710 García-Veigas etal.
(2013)
413.2 15.1 16.8 García-Veigas etal.
2013
489 14.9 16.5 0.708706 García-Veigas etal.
(2013)
514.6 0.708795 García-Veigas etal.
(2013)
529.6 15.4 17.5 0.708707 García-Veigas etal.
(2013)
Intermediate Sandstone Unit—ISU
553.7 14.8 16.4 This study
555 15.5 17.9 0.708771 This study
556.8 12.8 10 This study
559.5 15.6 18.1 This study
565.7 15.5 16.1 0.708710 This study
567.2 15.8 20.1 This study
569.5 16.4 20.3 This study
575.9 16 19.5 0.708779 This study
577 16.1 19.1 García-Veigas etal.
(2013)
581.4 16.4 20.9 This study
581.4 16.2 20 This study
584.2 15.8 18.7 This study
584.7 14.7 14.8 This study
586.8 15.9 20.1 0.708723 This study
586.9 0.708717 García-Veigas etal.
(2013)
Lower Halite Unit—LHU
617.6 13.6 16.3 García-Veigas etal.
(2013)
638 12.3 16 García-Veigas etal.
(2013)
648.1 0.708735 García-Veigas etal.
(2013)
659.5 11.9 17.1 García-Veigas etal.
(2013)
678.4 12.5 17.7 García-Veigas etal.
(2013)
703.5 15.1 17.3 García-Veigas etal.
(2013)
720 13.4 21 García-Veigas etal.
(2013)
724.1 12.8 21.7 0.708807 García-Veigas etal.
(2013)
Basal anhydrite bed
730 13.5 23.6 0.708869 García-Veigas etal.
(2013)
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The δ18OV-SMOW values from the ISU range from +12.8
to +16.4‰ with an average of +15.7 ± 0.5‰ (Fig.11).
The lowest value (+12.8‰) was found in a sample taken
at 556.8m deep. ISU samples have δ18OV-SM OW signatures
higher than those of Miocene marine sulfates (δ18O: +12
to +14‰; Claypool etal. 1980) and also higher than those
recorded from upper Triassic gypsum outcropping in the
Granada Basin (+12 ± 0.5‰; García-Veigas etal. 2013). The
values are also comparable to those reported for Tortonian
gypsum deposits from the Granada Basin (+16.8 ± 0.5‰;
Rouchy and Pierre 1979).
4.4.2 Strontium isotope ratios (87Sr/86Sr)
Strontium isotope analysis (87Sr/86Sr ratios) for several anhy-
drite nodules and halite samples from the ISU of the CMN-3
borehole are listed in Table1. The data are compared with
87Sr/86Sr values of Triassic dolomite (Gorzawski etal. 1989)
and Triassic evaporites (Ortí etal. 2014b) from the Betic
Cordillera, and from the Tortonian seawater (Hodell etal.
1991) in Fig.11.
The 87Sr/86Sr ratio ranges from 0.708710 to 0.708779
with an average of 0.708745, which is slightly higher com-
pared with the values reported for the ISU and UHU by
García-Veigas etal. (2013) (~ 0.708710; Table1). These
values are lower than the Tortonian seawater 87Sr/86Sr ratio
(~ 0.7089, Hodell etal. 1991) and higher than the 87Sr/86Sr
ratio measured in Triassic dolomite (~ 0.70825; Gorzawski
etal. 1989) and evaporite (0.707615–0.708114; Ortí etal.
2014b) samples from the Betic Cordillera.
5 Interpretations anddiscussion
5.1 Facies interpretation andenvironment
ofdeposition
In the ISU, six major lithofacies occur: (a) bioclastic sand-
stones (C1); (b) bioclastic siltstones (C2); (c) carbonate
mudstones (CB1); (d) carbonate/anhydrite interbeds (CE1);
(e) nodular anhydrite (E1), and (f) banded halite (E2).
The bioclastic sandstones (C1) have a clearly marine
provenance as testified to by the presence of abundant,
sometime well-preserved carbonate bioclasts, including
bryozoans, bivalves, coralline algae, and foraminifera.
This biota association is commonly found in temperate-
water, carbonate-platform deposits and is well exemplified
in sediments of this age (uppermost Tortonian) from the
nearby Almería basins (Sorbas and Vera basins and Cabo
de Gata area) (Martín etal. 1996, 2010; Betzler etal. 1997;
Braga etal. 2006; Puga-Bernabéu etal. 2007). Centimeter-
thick (3–20cm), sandstone layers are relatively well sorted,
exhibit a very irregular, clearly erosive base (Fig.5b), and
show a fining-upward grading (Fig.5a) and small (cm-
sized), wave-ripple and wave-ripple cross lamination on top
(Fig.5c). They are interpreted as storm layers, reworked by
very gentle waves, that were deposited in a very shallow
Fig. 10 A scatter plot of δ18O
and δ34S values of the evaporite
(sulfate) ISU of the Chime-
neas Halite Succession at the
CMN-3 borehole. Data are
completed for comparison with
values from the Granada Basin
(Chimeneas Halite Succession;
García-Veigas etal. 2013; sul-
fate isotope trend: green arrow)
and other Iberian Peninsula
Miocene basins (Calatayud,
Tajo, and Ebro basins; Utrilla
etal. 1992), along with values
for Miocene marine sulfates
(from Claypool etal. 1980;
Paytan etal. 1998), Upper Tri-
assic gypsum deposits (Keuper;
García-Veigas etal. 2013) and
Tortonian gypsum deposits
(Rouchy and Pierre 1979)
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barrier coastal lake. Similar deposits have been described by
Warren and Kendall (1985), Warren (1991), Evans (1995),
Kirkham (1997) and Strohmenger etal. (2011) in the Ara-
bian (Persian) Gulf. In the study example, terrigenous grains
are mainly dolomitic and metapelitic, high-grade metamor-
phic clasts indicating an Alpujárride source. Clast compo-
sition suggests they derived from some locally emerging
reliefs located to the S–SW of the salt depocenter (Fig.3),
at the position of the present-day Sierra de la Pera and Sierra
Tejeda where these types of rocks crop out extensively today
(Aldaya etal. 1980).
Bioclast remains in the siltstones (C2) indicate a marine
provenance as well. These sediments were also introduced
in the coastal lake during storms but, due to their small par-
ticle size, remained for a time in suspension before being
deposited.
Carbonate mudstone (CB1), calcitic in composition,
is considered to be the result of chemical precipitation
within the coastal lake. It may locally incorporate some
allochthonous, terrigenous/bioclastic silt grains as well as
some clay.
Anhydrite growths, closely associated with the carbon-
ate mudstones, occur in two ways, either as thin (< 1mm
to a few mm thick), anhydrite laminae intercalated within
carbonate mudstones (CE1 facies) (Fig.7c, d), or as layers
(up to 8cm thick) made up of a mosaic of tightly packed
nodules (E1 facies) (Fig.8a). Anhydrite laminae (in CE1
facies) formed by the replacement of thin, selenite-gypsum
levels. As stated above, they show ghost relics of presum-
ably former micro-selenite gypsum crystals (Fig.7e, f).
Anhydrite nodules (E1 facies) exhibit a subtle, chicken-wire
internal structure. This sort of structure has been described
in modern marine and continental evaporitic sabkha, where
anhydrite, crystallizing from interstitial hypersaline waters
inside the mud, results from synsedimentary transformation
after gypsum (e.g. Warren 2006, 2016). High temperatures
(~ 35–45°C) at shallow (1–2m) depths can lead to the con-
version from gypsum to anhydrite in the near-surface when
Fig. 11 a Condensed geochemical evolution of the Chimeneas Hal-
ite (CMN-3 borehole). Sulfate isotope composition (δ18O and δ34S)
and 87Sr/86Sr ratio data come from García-Veigas etal. (2013) and
from this work (ISU interval). b Vertical evolution of geochemical
profiles (δ34S, δ18O and 87Sr/86Sr) of the ISU of the Chimeneas Hal-
ite Succession at the CMN-3 borehole. Sulfate isotope values for the
ISU (this work) are compared with those from Miocene marine sul-
fates (δ18O from Claypool etal. 1980, δ34S from Paytan etal. 1998),
Upper Triassic (Keuper) gypsum deposits (García-Veigas etal. 2013)
and Tortonian gypsum deposits from the Granada Basin (Rouchy
and Pierre 1979). Strontium isotope ratios (this work) are compared
with 87Sr/86Sr values of Triassic dolomite (Gorzawski etal. 1989) and
Triassic evaporites (Ortí etal. 2014b) from the Betic Cordillera, and
Tortonian seawater (Hodell etal. 1991)
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pore-fluid salinity is close to halite saturation (Warren 2016).
According to Holser (1979), in evaporitic settings dominated
by gypsum and halite, anhydrite replacement can occur at
25°C. Such temperatures are common on modern evaporitic
playas and could have been reached in the Granada Basin in
the latest Tortonian. Anhydrite-nodule formation took place
within the carbonate mud in the saline lake by the replace-
ment of previously formed gypsum crystals.
Primary halite (in E2 facies) consists of centimeter- to
decimeter-thick strata of massive, white (cloudy) hal-
ite, exhibiting a chevron structure (Fig.8c) and contain-
ing abundant, sometime aligned, fluid inclusions (Fig.8d,
e). Halite crystals are up to 1cm in size. This halite was
deposited at times when the coastal lake turns into a fully
saline lake. Lenticular to prismatic gypsum crystals occur-
ring at the base of some halite layers are interpreted as hav-
ing precipitated during the first stages of concentration in the
saline lake. Chevron fabric in salt is interpreted as forming
by competitive growth of bottom-nucleated (subaqueous)
crystals (e.g. Lowenstein and Hardie 1985). Relict zones
of chevron growth fabrics outlined by brine inclusions are
the most common primary texture preserved in modern and
ancient salt deposits (e.g. Hardie 1984; Lowenstein and Har-
die 1985; Schléder and Urai 2005).
From facies interpretation it is inferred that the deposition
environment for the whole ISU was that of a coastal lake,
Fig. 12 Major stages in the temporal evolution of the coastal lake
during deposition of the ISU. a Synsedimentary formation of the nod-
ular anhydrite beds (after gypsum) within carbonate mud in a very
shallow, lacustrine evaporitic environment. b Persistent marine influ-
ence and frequent inundation of the lake during major storms with
deposition of bioclastic sands and silts. c Perennial saline lake with
precipitation of micro-selenite gypsum and halite. Sporadic inunda-
tions resulted in the partial dissolution of some of the halite layers
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more likely isolated from the open sea by a sand barrier,
subjected to strong salinity fluctuations.
5.2 Coastal-lake evolution
As pointed out above, vertical facies distribution in the ISU
sequence shows three distinct intervals (Fig.4). At the bot-
tom of the sequence (first interval) nodular anhydrite layers
are frequently found. In the middle (second interval) bioclas-
tic sediments are dominant. At the top (third interval) halite
layers and carbonate/anhydrite interbeds are ubiquitous.
These intervals in turn mark three different stages in the
temporal evolution of the lake (Fig.12). Syn-depositional
anhydrite nodules formed close to surface in a very shallow,
lacustrine evaporite environment (stage 1, Fig.12a). Stage 1
is considered to be the shallowest in the lake’s evolution. The
environmental change was then to a deeper, perennial lake
undergoing frequent marine flooding storm events resulting
in bioclastic deposition (stage 2, Fig.12b). A shallower, per-
ennial saline lake was finally established (stage 3, Fig.12c),
with major evaporite (anhydrite and halite) formation. Anhy-
drite laminae are thought to be the result of replacement
of bottom-nucleated (subaqueous), micro-selenite gypsum.
Chevron halite beds resulted from the precipitation and com-
petitive growth of bottom-nucleated, salt crystals.
Marine flooding events resulted in partial dissolution of
the primary halite crust. Dissolution concentrated at first at
grain boundaries (e.g. Shearman 1970), creating later larger,
elongated cavities. Dissolution of the saline crust was fol-
lowed by the deposition of a thin, carbonate-mud lamina
(Fig.13), precipitated from the brackish water resulting from
the incorporation of under-saturated water immediately after
the flooding event.
The Granada Basin saline lake thus experienced repeated
periods of flooding (and partial dissolution of evaporites),
and of salt crystallization (Fig.12c). Syn-depositional tex-
tures recorded in recent salt-pans (e.g. Shearman 1970;
Lowenstein and Hardie 1985; Casas and Lowenstein 1989),
which have also been reported in ancient salt deposits (e.g.
Wardlaw and Schwerdtner 1966; Casas and Lowenstein
1989; Benison and Goldstein 2001; Schléder and Urai 2005),
are consistent with our interpretations for the ISU. Salt-pan
evolution normally shows repeated periods of flooding,
evaporative concentration and, sometimes, desiccation. The
flooding stage incorporates unsaturated floodwaters into the
salt-pan, converting it into a temporary shallow brackish
lake. As reported by many authors, both fresh- and marine-
water flooding events are frequent in marginal-marine, saline
pans (e.g. Lowenstein and Hardie 1985; Schreiber and El
Tabakh 2000; Benison and Goldstein 2001; Warren 2006,
2010, 2016). Fresh waters typically come from rainstorm
runoff (and snow meltwater), while seawaters are washed
onto the pan by storm floods and, rarely, at high spring tides.
In our example, the absence of desiccation structures (mud
cracks, tepees, etc.) indicate that ‘the saline pan cycle’ in the
sense of Lowenstein and Hardie (1985) was never completed
up to the point of full desiccation of the coastal, saline lake.
5.3 Diagenetic evolution
The petrographic study shows the effect of several diagenetic
processes, taking place at different times (stages), after sedi-
ment deposition, in the course of the diagenesis (Fig.14).
5.3.1 Syn-depositional diagenetic stage
As discussed above, anhydrite is replacing former gypsum
crystals in nodules (E1 facies) and layers (CE1 facies) within
the carbonate mudstone. These anhydrite replacements are
thought to have occurred very early, within the deposition
environment. Dissolution surfaces developed on top of halite
layers after flooding events. Brackish waters were respon-
sible for this dissolution (dissolution stage 1, herein). Dis-
solution cavities also developed underneath, inside the salt
(Figs.13, 15a). Clear halite, precipitated as void-filling
cement inside primary-halite, dissolution cavities, exhibits
no preferred crystal orientation, growth direction or elonga-
tion. This halite, which often truncates the chevron halite
crystals, is interpreted as newly formed, synsedimentary
cement precipitated from interstitial groundwater brines
(e.g. Shearman 1970; Lowenstein and Hardie 1985; Casas
and Lowenstein 1989). This limpid, clear halite also occurs
Fig. 13 Core slab showing dissolution on the topmost part of a halite
layer (H). Dissolution was followed by the deposition of a thin, car-
bonate-mud lamina (M, yellow arrow), which occurs now lining the
former dissolution surface. White arrow points to a small anhydrite
(A) nodule
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530 Journal of Iberian Geology (2018) 44:513–537
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as intra- and inter-granular pore-filling cement (Fig.5e, i)
in underlying clastic facies.
5.3.2 Early diagenetic stage
Early diagenesis is reflected in the halite deposits (E2 facies)
by re-crystallization processes, back reactions, and replace-
ments. The main early diagenetic process is halite re-crys-
tallisation of pre-existing halite (depositional primary halite
and syn-depositional halite cement) (Fig.8b, c), as a result
of repetitive dissolution (dissolution stage 2)-precipitation
cycles (Fig.15b). Recrystallized halite shows up as an equi-
granular mosaic, in which the curved crystal boundaries
meet at triple junctions with angles of approximately 120°
(Fig.8g). The fact that this mosaic-type halite exhibits disso-
lution gulfs like those described in the primary, depositional
halite, and that it sometimes fills in intergranular pores in
the immediately underlying sandy and silty layers (C1 and
C2 facies), points to an early diagenetic process occurring
close to surface. According to Lowenstein and Hardie (1985)
and to Warren (2016), dissolution and re-crystallization of
primary halite into a secondary halite occurs under a few
cm to meters of burial depth from brine reflux following
the dissolution of the primary halite. In our example the
reflux brines seem to have moved preferentially along bed-
ding planes since re-crystallization locates preferentially at
the base and on top of the halite layers, resulting in their final
banding display. Early re-crystallization processes derived
from repetitive dissolution–precipitation processes have
been reported in desiccated halite-floored salinas and salt
lakes (e.g. Handford 1982).
Halite replacement of gypsum is shown by gypsum pseu-
domorphs keeping external morphologies, internal crystal
structures and associated impurities (Figs.9a, b, 15b). This
replacement apparently took place early within a few tens
of centimeters of burial depth. According to Schreiber and
Walker (1992) and Hovorka (1992), halite replacement of
gypsum is linked to temperature variations and associated
changes in the solubility of both minerals following chemi-
cal changes in the brine composition, in our case probably
related to salinity changes of the water in the repetitive,
early-diagenetic, dissolution–precipitation cycles.
Nodular anhydrite replaces pre-existing halite crystals
(Fig.8h) and locates at the base and on top of halite beds (E2
facies; Fig.7a). The 87Sr/86Sr and the δ34SV-CDT signatures
of halite and anhydrite are quite similar, pointing to an early
process preventing a strong isotopic fractionation.
Pyrite is found as framboid growths (spherical aggregates
of micron-sized pyrite crystals; Love and Amstutz 1966;
Sweeney and Kaplan 1973). Framboid pyrite is associated
with plant debris (Figs.6f, 9c). According to Berner (1970),
it results from the reaction of sulfide (produced via bacterial
sulfate reduction; BSR) with either Fe3+ (in sediments) or
Fe2+ produced by bacterial Fe3+ reduction (Lovley 1991).
Thus, the presence of pyrite suggests that BSR processes
occurred during early diagenesis in the ISU sediments.
5.3.3 Intermediate- tolate-diagenetic stages
Authigenic silica, equigranular quartz (megaquartz), and
fibrous quartz (chalcedony and lutecite), replaces pre-exist-
ing evaporites (i.e. secondary, mosaic halite, and anhydrite
crystals). The sequence of silica replacement is from mega-
quartz (Fig.9d, e), to chalcedony (Fig.9f), and to lutecite
(Fig.9g–i). All these quartz varieties have been reported as
a testimony of vanished evaporites (Folk and Pittman 1971;
Fig. 14 ISU sequence of diage-
netic events (CMN-3 borehole)
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Friedman and Shukla 1980). According to Folk and Pittman
(1971) and Arbey (1980) silica-crystal size and morphology
are controlled by the sulfate content within the diagenetic
environment. Megaquartz forms in non-sulfated (or poorly
sulfated) diagenetic environments. On the contrary, chal-
cedony and lutecite crystallize in sulfated diagenetic envi-
ronments. A low pH promotes chalcedony formation, while
a high pH favors lutecite crystallization (Arbey 1980).
According to Tucker (1976), Geeslin and Chafetz (1982),
and Maliva (1987), anhydrite replacement by quartz is an
early diagenetic process. However, silicification of anhydrite
nodules has been related to burial diagenesis when hydro-
carbon inclusions are found inside the megaquartz (Ulmer-
Scholle etal. 1993), and hydrothermalism cannot be ruled
out to explain the crystallization of lutecite (Arbey 1980).
According to García-Veigas etal. (2013), halite precipita-
tion in the UHU of the Chimeneas Halite is associated with
significant, underground water inputs, including CaCl2-rich,
sulfate-depleted, hydrothermal waters. Percolating brines
from the UHU with such as fluid chemistry could explain
the formation of megaquartz in the underlying ISU.
The diagenetic conditions in the ISU appear to have
changed from non-sulfated (or poorly sulfated), as con-
firmed by the anhydrite replacement to megaquartz, to more
sulfated conditions, as attested to by the crystallization of
chalcedony and lutecite. In this respect, it is worth noting
that the Chimeneas Halite Unit is capped by a thick gypsum
sequence (Fig.1) reflecting a major shift in brine composi-
tion to sulfate-rich fluids, and presumably, in the same way,
in the fluids percolating through the underlying deposits (i.e.
the ISU deposits).
Celestine occurs as a late-diagenetic replacement of lute-
cite after mosaic halite (Fig.9i). Celestine occurrences are
commonly interpreted as result of the diagenetic interac-
tion of Ca-sulfates with Sr2+-rich groundwater (e.g. War-
ren 2016). The dissolution of gypsum or anhydrite supplies
the SO4, which combined with Sr2+-rich waters lead to the
precipitation of celestine. The large celestine orebodies
(Montevive and Escúzar ore deposits; Martín etal. 1984;
García-Veigas etal. 2015) of the Granada Basin (Fig.1)
formed by early diagenetic replacement of gypsified stro-
matolites (García-Veigas etal. 2015). According to García-
Veigas etal. (2015), dissolved gypsum was the main sulfate
source, while marine waters and diagenetic–hydrothermal
CaCl2 brines were the main strontium sources. The small
celestine crystals found at ISU (Fig.10i) could be the result
of recycling by underground waters of some of the previ-
ously formed celestine.
5.4 Brine parenthood
Basin architecture, tectonic evolution and climate of the
Granada Basin in the latest Tortonian were appropriate to
Fig. 15 a Simplified diagram illustrating syn-depositional overprints
on primary, chevron-type halite (after Shearman 1970, modified).
The textural evolution of halite crust shows: (1) dissolution on top of
primary halite, (2) deposition of mud, and (3) precipitation of halite
cement in the dissolution cavities. b Simplified diagram illustrating
re-crystallization of halite and formation of secondary, mosaic-halite
driven by repetitive dissolution/precipitation (evaporative concentra-
tion) cycles
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532 Journal of Iberian Geology (2018) 44:513–537
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establish a basin-wide, evaporite depositional environment
(Martín etal. 1984) (Fig.3). In the Granada Basin salt suc-
cession, three thick halite-bearing units were identified: the
‘Lower Halite Unit’ (LHU), the ‘Intermediate Sandstone
Unit’ (ISU), and the ‘Upper Halite Unit’ (UHU). Geochemi-
cal markers (brine inclusion composition, bromine content,
and isotope data) indicate a marine origin for the LHU,
with a growing influence of non-marine input higher in the
sequence. The environment of deposition for the LHU was
envisaged as a salt-concentrated, marine lagoon (García-
Veigas etal. 2013). In the ISU and UHU, the non-marine
influences became progressively more significant, with geo-
chemical markers suggesting recycling of previously formed
halite and also contribution from hydrothermal waters. The
UHU environment of deposition was that of a continental
salt-pan, clearly disconnected from the sea (García-Veigas
etal. 2013).
The ISU marks the transition from a marine to a conti-
nental setting for the precipitation of the salt in the Granada
Basin. As stated above, the deposition environment for the
ISU was that of a coastal lake, isolated from the open sea
by a sand barrier and subjected to strong salinity fluctua-
tions. In the ISU evaporitic deposits, 87Sr/86Sr isotopic val-
ues (Fig.11) are slightly lower than those of the Tortonian
seawater (Hodell etal. 1991). The strontium (87Sr/86Sr) and
sulfur (δ34S) isotope compositions are midway between the
marine Miocene (Paytan etal. 1998; Hodell etal. 1991)
and the Triassic basement (evaporite and dolomite) values
(Gorzawski etal. 1989; García-Veigas etal. 2013; Ortí etal.
2014b) (Fig.11). All this suggests a mixed origin, between
non-marine (underground/hydrothermal) and marine waters,
for the parenthood brines.
ISU sulfate isotope values (δ34S and δ18O) are very close
to those found by Rouchy and Pierre (1979) for the Granada
Basin Tortonian gypsum (Figs.10, 11). The sulfur and oxy-
gen signatures are also similar to some of those recorded in
continental evaporites from the Calatayud and Tajo Basins
(Utrilla etal. 1992; Fig.10). According to these authors,
such signatures resulted from the remobilization of marine
Triassic evaporites. Sulfur and oxygen enrichment in the
ISU evaporites compared to Triassic evaporites (Figs.10,
11) could be the result of the combined effects of bacte-
rial sulfate reduction (BSR), fractional crystallization, and
strong evaporation. According to Antler etal. (2013) a posi-
tive covariance between oxygen and sulfur isotope compo-
sitions points to BSR. During BSR the sulfur and oxygen
isotope values of the residual, dissolved sulfate, and sulfide
increases and decreases, respectively (Seal etal. 2000). Our
data are in line with such considerations and suggest strong
BSR in an anoxic environment during ISU deposition.
Additional contribution of CaCl2-rich hydrothermal flu-
ids to the ISU brines, as suggested by García-Veigas etal.
(2013), cannot be disregarded. Diagenetic-hydrothermal,
CaCl2-rich brines have been proposed to explain the
‘selected removal’ of sulfate ions in many evaporite basins
(Ayora etal. 1994b, 2001; Cendón etal. 2008; García-Veigas
etal. 2009; Lowenstein and Risacher 2009; García-Veigas
etal. 2013; 2015), favoring the precipitation of halite instead
of gypsum. CaCl2-rich brines form by diagenetic interac-
tions between rock/sediment and heated groundwaters (e.g.
Hardie 1990; Lowenstein and Risacher 2009). These brines
can reach the subsurface by convective- or topographically
driven circulation (Hardie 1990), associated mainly with
volcanism and faulting.
The eastern Betic Cordillera, as revealed by many
authors, recorded the synchronous interaction of tectonic,
volcanic, hydrothermal and evaporitic processes in the
course of the late Miocene (e.g. Benito etal. 1999; Dinarès-
Turell etal. 1999; Kuiper etal. 2006; Playà and Gimeno
2006). However, such as evidence does not exist, particularly
concerning volcanism, in the case of the Granada Basin.
Basinal brines migrating upwards along faults in a tectoni-
cally active basin seem to be the most plausible scenario
for the Granada Basin. As stated above, major fault systems
in the Granada Basin have NW–SE and E–W orientations,
such as the ‘Cádiz-Alicante fault’ (Sanz de Galdeano 2008),
delineating its northern margin (Fig.2), and the ‘La Malahá
fault’ (Rosino 2008), limiting the salt depocenter (Fig.3).
The reactivation of such faults at the end of the Miocene
(Montenat etal. 1990; Sanz de Galdeano 1990; Rodríguez-
Fernández and Sanz de Galdeano 2006) promoted convec-
tive flow of groundwaters in areas of the basin with elevated
heat flows and could also have facilitated the circulation of
hydrothermal fluids (i.e. CaCl2-rich brines) presumably
coming from deep underground (Fig.16).
The Lorca basin is a comparable example to that of the
Granada basin (García-Veigas etal. 2013). This basin is also
limited by the Cádiz-Alicante fault (Sanz de Galdeano 1990,
Fig.2a), and its Upper Miocene record also encompasses a
marine, an evaporitic, and a continental succession, from
bottom to top (Geel 1976; Krijgsman etal. 2000; Taberner
Fig. 16 Hydrology model during the deposition of the ISU in the
Granada Basin
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533Journal of Iberian Geology (2018) 44:513–537
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etal. 2000). The marine sequence culminated with the depo-
sition of a ‘Gypsum Member’, which also includes a thick
(> 200m) halite deposit (García-Veigas 1993; Ayora etal.
1994b; Taberner etal. 2000). Based on geochemical evi-
dence, García-Veigas (1993) and Ayora etal. (1994b) dif-
ferentiated within the Halite Unit a marine Lower Part and
a continental Upper Part. According to these authors, upper
halite deposition took place under non-marine conditions
and resulted from the input into the basin of saline under-
ground waters recycling older evaporites (i.e. Keuper evap-
orites) and hydrothermal, CaCl2-rich waters coming from
deep underground through major faults. Sedimentological
interpretations combined with geochemical analyses in both
Granada and Lorca basins point to a similar evolution from
marine to non-marine conditions. Recent biostratigraphic
studies (Corbí etal. 2012), however, suggest that the evapo-
rites of the Granada Basin could be younger than those from
the Lorca basin.
6 Conclusions
The Granada Basin contains a thick salt sequence (Chime-
neas Halite) formed at the time the basin desiccated during
the latest Tortonian. The salt succession comprises three
halite-bearing units: the ‘Lower Halite Unit’ (LHU), the
‘Intermediate Sandstone Unit’ (ISU), and the ‘Upper Halite
Unit’ (UHU). Detailed sedimentological and petrographical
core observations, combined with isotopic analysis (δ34S and
δ18O in sulfates and 87Sr/86Sr in sulfate and halite samples),
of the ISU yield the following conclusions:
In the ISU deposits, six major facies-types are distin-
guished: bioclastic sandstones (C1); bioclastic siltstones
(C2); carbonate mudstones (CB1); carbonate-anhydrite
interbeds (CE1); nodular-anhydrite beds (E1) and banded
halite (E2). The bioclastic sandstones (C1 facies) are inter-
preted as marine sediments redeposited, as storm layers,
inside a coastal lake. The bioclastic siltstones (C2 facies)
are considered to be suspension deposits, of marine ori-
gin, introduced as well in the lake by storms. Carbonate
mudstones (CB1 facies) are thought to represent inorganic,
lacustrine chemical precipitates. Anhydrite laminae (in CE1
facies) are interpreted to have formed at the bottom of the
lake by syn-sedimentary replacement of micro-selenite-gyp-
sum. Nodular-anhydrite (E1 facies) seems to have developed
within carbonate mud by the replacement of gypsum, and
primary-halite crystals (in E2 facies) are considered to be
insitu growths, deposited sub-aqueously at the bottom of
the lake.
The environment of deposition for the ISU was that of a
coastal lake, isolated from the open sea by a sand barrier. We
propose the precise boundary of the onset of marine to non-
marine conditions in the Granada Basin to be placed during
the ISU deposition. Three main stages are recognized in the
evolution of the lake. At stage 1, syn-depositional anhy-
drite nodules formed close to the surface in a very shallow
lacustrine environment. At stage 2, frequent marine flood-
ing storm events resulted in significant bioclastic, sandy
sedimentation in a presumably deeper lake. At stage 3, a
shallower, perennial saline lake was established and major
evaporite deposition (anhydrite after micro-selenite gypsum,
and primary chevron halite) took place. Dissolution surfaces
developed on top of halite layers (E2 facies) after flooding
events. Clear-halite, void-filling cement precipitated syn-
depositionally inside primary-halite, dissolution cavities.
Early diagenesis is shown by: halite re-crystallization of
pre-existing halite (depositional halite and syn-depositional
halite cement), as a result of repetitive dissolution–pre-
cipitation cycles; gypsum replacement by halite, and halite
replacement by nodular anhydrite. The presence of fram-
boid pyrite associated with plant debris suggests BSR taking
place during early diagenesis in the ISU sediments.
Intermediate- to late-diagenetic processes are silica
(megaquartz, chalcedony, and lutecite) replacement of halite
and anhydrite, and celestine formation. Megaquartz formed
within a poorly sulfated, diagenetic environment. Percolat-
ing brines, at the time of deposition of the UHU, seem to
be the most likely fluids responsible for the existence of
megaquartz in the underlying ISU. Chalcedony and lutecite
crystallized sometime later, presumably during deposition
of the gypsum sequence occurring on top of the Chimeneas
Halite Unit. This implied a major change to a sulfate-rich
brine and, in a similar way, in the chemistry of the fluids
percolating through the underlying ISU deposits. Celes-
tine found at ISU replacing lutecite could be the result of
the interaction with Sr-rich underground waters dissolving
some of the previously formed celestine (Montevive/Escúzar
Celestine ore deposits).
Strontium and sulfur isotope compositions are midway
between the Miocene marine signature and the Triassic
basement, evaporite and dolomite signatures. This sug-
gests a mixed origin, between non-marine (underground)
and marine waters, for the parenthood brines, during evap-
orite deposition. Additional contribution of hydrothermal,
CaCl2-rich basinal brines, migrating upwards along faults,
cannot be ruled out.
Acknowledgements This work is part of the ‘Study MFT08.003’ (IFP
Énergies Nouvelles, France). The initial part of this work was car-
ried out at IFP Énergies Nouvelles laboratories in Rueil-Malmaison
(France), during an internship stay of the senior author (ALQ) spon-
sored by IFP. JMM work was funded by the project ‘Productores de
carbonato en plataformas carbonatadas neógenas de la Cordillera
Bética. Factores que controlan la composición y la resedimentación’
(CGL2013-47236-P) (2014-18) (Ministerio de Economía y Competi-
tividad, Spain and Fondo Europeo de Desarrollo Regional FEDER).
Thanks are given to Dr. Javier García-Veigas (University of Barcelona)
for the critical reading of an earlier version of this paper and for helping
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534 Journal of Iberian Geology (2018) 44:513–537
1 3
with the isotopic analysis carried out at the CCiT-UB (Centres Cientí-
fics i Tecnològics, Universitat de Barcelona), and to SUERC (Scottish
Universities Environmental Research Centre) for the strontium isotopic
analysis. We would also thank Dr. Ángel Puga-Bernabéu and Dr. Julio
Aguirre (University of Granada) for their helpful contribution to this
work. The authors would like to express their sincere thanks to the Edi-
torial Board and anonymous reviewer for their suggestions to improve
the paper. We thank David Nesbit for the correction of the English text.
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Aliations
AdriánLópez-Quirós1 · MickaelBarbier2· JoséM.Martín3· XavierGuichet2
* Adrián López-Quirós
alquiros@iact.ugr-csic.es
1 Instituto Andaluz de Ciencias de la Tierra, CSIC-
Universidad de Granada, Avda. las Palmeras 4,
18100Armilla,Granada, Spain
2 IFP Énergies Nouvelles, 1-4 Avenue de Bois-Préau,
92852Rueil-Malmaison, France
3 Departamento de Estratigrafía y Paleontología, Universidad
de Granada, Campus de Fuentenueva s.n., 18002Granada,
Spain
Author's personal copy