Trabajos geoquimica by Cesar Menor-Salvan

Research Papers on Geochemistry

César Menor Salván May 2015

Summary The present volume comprehends the research papers published by Dr. Cesar Menor Salvan regarding Organic Geochemistry and the search of life biosignatures in the geological record. Dr. Menor-Salvan implemented a new research line on molecular fossils, Organic Geochemistry and relation life-mineral deposits at the Centro de Astrobiologia (CSIC-INTA) in 2009. This research line was blocked by the science cuts in 2012.

Papers published as main contributor or principal researcher

Menor-Salvรกn C, Simoneit BRT, Ruiz-Bermejo M, Alonso J., Huiming L (2015). Structural identification of 1,6-dimethyl-5-alkyltetralins, a new labdane biomarker family from Cretaceous ambers. Organic Geochemistry; Sent. Menor-Salvรกn C, Najarro M, Velasco F, Rosales I, Tornos F, Simoneit BRT (2010) Biological diterpenes preserved in Lower Cretaceous amber from Basque-Cantabrian basin (El Soplao, Cantabria, Spain). Paleochemotaxonomical aspects. Organic Geochemistry, 41: 10891103.DOI: 10.1016/j.orggeochem.2010.06.013 Menor-Salvรกn C, Tornos F, Fernรกndez Remolar D, Amils R (2010) Association between catastrophic paleovegetation changes during Devonian-Carboniferous boundary and the formation of giant massive sulfide deposits. Earth and Planetary Science Letters, 299: 398408. DOI: 10.1016/j.epsl.2010.09.020

Journal of Nuclear Materials 462 (2015) 296–303

Contents lists available at ScienceDirect

Journal of Nuclear Materials journal homepage: www.elsevier.com/locate/jnucmat

Study of the alteration products of a natural uraninite by Raman spectroscopy L.J. Bonales a,⇑, C. Menor-Salván b, J. Cobos a a b

Centro de Investigaciones Energéticas, Medioambientales y Tecnológicas, CIEMAT Avenida Complutense, 40, 28040 Madrid, Spain Centro de Astrobiología (CSIC-INTA), Ctra. Torrejón-Ajalvir, km 4, 28850 Torrejon de Ardoz, Spain

a r t i c l e

i n f o

Article history: Received 7 October 2014 Accepted 10 April 2015 Available online 17 April 2015

a b s t r a c t Uraninite is a mineral considered as an analogue of the spent fuel, and the study of its alteration products has been used to predict the secondary phases produced during the fuel storage under specific environmental conditions. In this work, we study by Raman spectroscopy the alteration by weathering of the primary uraninite from the uranium deposit of Sierra Albarrana. The identification of the different secondary phases is based on the analysis of the symmetrical stretching vibration of the uranyl group (UO2+ 2 ), which allows the identification of individual uranyl phases and can be used as a fingerprint. Additionally, we show in this work a new approach to perform a semi-quantitative analysis of these uranium minerals by means of Raman spectroscopy. From this analysis we found the next sequence of alteration products: rutherfordine in contact with the uraninite core, then a mixture of uranyl silicates: soddyite, uranophane alpha and kasolite. Soddyite prevails in the inner part while uranophane alpha predominates in the outer part of the sample, and kasolite appears intermittently (1.0–3.3 mm; 4.6–7.1 mm and 8.8–10 mm). Ó 2015 Published by Elsevier B.V.

1. Introduction High-level nuclear waste, such as irradiated UO2 (spent fuel) will be disposed in an underground repository. It is expected that the spent fuel will be exposed to groundwater after storage times of the order of some thousand years, when the containers surrounding the waste may be breached. Identiﬁcation of the reaction products generated by the interaction of the waste form with water is required to characterize the repository performance [1]. All scenarios describing the spent fuel–groundwater contact require extrapolations to the far future of a complex system, whose components are not all well-deﬁned. It is expected that the conditions in a spent fuel deep geological disposal will be reducing. Nevertheless, the amount of oxidizing species near the spent fuel surface will increase due to the radiolysis of water caused by the ionizing radiation associated with the fuel [2], which is dominated at the predicted time of the breached containers by the a-decay [3,4]. In particular, only those radiolysis products, which are formed in the water layer near the fuel surface, i.e. within <50 lm of the fuel surface, are effective in causing the fuel oxidation [5]. Therefore, both dissolution and precipitation processes ⇑ Corresponding author at: Centro de Investigaciones Energéticas, Medioambientales y Tecnológicas, CIEMAT, Departamento de Energía, Unidad de Residuos de Alta Actividad, Av Complutense, 40, 28040 Madrid, Spain. Tel.: +34 913462576. E-mail address: laura.jimenez@ciemat.es (L.J. Bonales). http://dx.doi.org/10.1016/j.jnucmat.2015.04.017 0022-3115/Ó 2015 Published by Elsevier B.V.

under this conditions will affect the overall behavior of the fuel matrix. Depending of the surface/volume ratio, secondary phases will appear on the spent fuel surface as alteration products [6]. Different approaches can be done in order to study the dissolution/precipitation processes of spent nuclear fuel (SF) and understand the potential migration of uranium under repository conditions during millions of years. On one hand, the use of mathematical models [7,8] allows predicts SF behavior in the long-term at the expected conditions. These theoretical methods require the knowledge of physico-chemical parameters. These must be obtained experimentally in laboratory assays for SF analogues, such as uranium dioxide UO2 [9,10] or SIMFUEL [11,12]. On the other hand, the studies of natural analogues have been very successful, to understand different aspects of the SF corrosion processes at longer storage times [13–24]. The uraninite is known as the natural analogue of the SF, and studies about its dissolution [13,14] and its corrosion [15,16] at different conditions have been performed for decades. The consideration of this mineral as an analogue of the SF is mainly due to two reasons: (1) uraninite is a non-stoichiometric compound with a chemical composition very similar to that of spent fuel (UO2 > 95%) [17,18], although natural uraninite contains variable quantities of radiogenic lead which could generate secondary U– Pb phases after alteration; (2) both materials have a cubic ﬂuorite structure, (space group Fm3m). Studies of different uraninites have been widely performed. Some studies [19–22] have shown the

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temporal sequence of the alteration products of natural uraninites at different geochemical conditions. The general trend of this sequence was previously recognized by Frondel [23,24] and it is still widely accepted as: (1) uranium oxides, (2) uranyl oxyhydroxides and (3) uranyl silicates; and the specific alteration products depend on local conditions. Characterization of the alteration products of natural uraninites has been conventionally done by the combination of different techniques: optical petrography (OP) [19], scanning electron microscopy/energy-dispersive X-ray analysis (SEM/EDS) [19–22] and electron microprobe analysis (EMPA) [19–21]. There are nowadays other techniques that can provide a complete information by themselves, without using complementary techniques, such as Raman spectroscopy. The advantages of the Raman spectroscopy are as follow: (1) the samples do not require any special preparation, so possible alterations due to these previous steps are avoided by using this technique; (2) this technique allows the analysis of very small samples with little background interference [25–27]. However, knowledge of Raman spectra of uranyl-based minerals is still rather limited except the effort of Frost et al. and Amme et al. [28,36,37]. This work focuses on the complete identification by Raman spectroscopy of the supergenic uranyl phases produced by alteration of a natural analogue of the spent fuel: the uraninite from Sierra Albarrana (Spain). To the best of our knowledge this is the first study of uraninite and its associated secondary uranium minerals from the uranium deposit of Sierra Albarrana, and it has been shown that the complete characterization of uranium secondary phases is possible using only one characterization technique. Emphasis on the development of a new approach, the semi-quantitative analysis of these uranium minerals by Raman spectroscopy has been done.

2. Materials and methods 2.1. Mineral sample and geological setting The sample studied is a ‘‘uraninite + gummite’’ from Sierra Albarrana (Córdoba, Spain), kindly provided by the Museum of Natural Sciences of Alava (Vitoria, Spain). It has been collected during the uranium extractive activity in 1960. The sample structure corresponds to the ideal gummite occurrence [23,24] (see Fig. 1a): a veined central core black to brownish black and yellow to orange or greenish yellow surrounding zone, vitreous to dull or earthy, formed by several supergenic minerals. The uranium-rare earth mineralization at Sierra Albarrana (Cordoba, Spain) [29] is distributed in a complex pegmatite field of granitic composition. The pegmatites are not related with plutonic bodies and are syn-metamorphic and formed by anatexis associated to medium to high grade metamorphism during Variscan. The pegmatite field is hosted by the Cambrian quartzite and gneiss of the metamorphic nucleus of the Albarrana formation [30]. The pegmatite forms irregular bodies and veins parallel to the Variscan structures. The pegmatite mineralogy is controlled by the metamorphic grade and the accessory minerals include uraninite, thorite, brannerite, beryl, schorl , rutile, ilmenite, allanite–(Ce), zircon, monazite–(Ce), xenotime, columbite–tantalite, chrysoberyl and Fe–Mn–Mg–Caphosphates [31,32]. From a geochemical point of view, the uraninite-gummite samples studied in this work belong to a pegmatite of the Muscovite-rare element class in the Cherny classification [33] and the type ‘‘Dieresis’’ in the classification of Gonzalez del Tánago [34], hosted by rich biotitic-muscovitic gneisses. This type of pegmatite was worked for the extraction of uranium.

297

2.2. Preparation of the sample The sample studied in this work was obtained from the Mineralogy collection of the Natural Sciences Museum of Alava (MCNA, Vitoria, Spain). The sample is from the Dieresis uranium mine (Sierra Albarrana, Córdoba, Spain). The sample was cut using a diamond saw and polished. The thick polished sections obtained were subjected to spectroscopic analysis. 2.3. SEM–EDS A polished section of the sample was analyzed under a Jeol 5600-LV scanning electron microscope equipped with an Oxford Industries INCA X-sight energy dispersive X-ray spectrometer. Backscattered electron images and energy dispersive spectra were obtained on the sample mounted on Al stubs and without coating (V = 20 kV I = 85 lA, electron beam diameter 1 lm). 2.4. Raman spectroscopy technique The Raman spectroscopy was carried out using a Horiba LabRam HR evolution spectrometer (Jobin Yvon Technology). A red laser of HeNe with a wavelength of 632.81 nm and an operation power of 20 mW was used as the excitation source. The laser was focused onto the sample using 20 objective at the confocal microscope BX4 with confocal 800 mm; the scatter light was collected with the same objective and then dispersed with a JobinYvon spectrometer (600 gooves/mm), and detected with a peltier cooled CCD detector (256 1024 pix.). The spectral resolution was about 1 cm 1 per pixel. 2.5. Raman mapping procedure The surface of the sample was analyzed by acquiring 100 spectra in different points separated 100 lm from each other. The ﬁrst one corresponds to the center of the sample and the rest were located on a line going from the center of the sample outwards, (see Fig. 1a and b). In Fig. 1 we have indicated the position at the sample in which the spectra were taken. As it can be seen, we have divided the sample in eight regions for further analysis. The choice of the different region lengths was performed by visual analysis. Regions 1, 2 and 3 correspond to zones of different darkness near to the core of the sample, whereas zone 4, 5, 6, 7 and 8 correspond to different yellow tones of the sample, and are separated by veins. The protocol used is a combination of the line-mapping and step-by-step procedures, as described below: The sample is placed in the motorized x–y stage under the microscope objective and focused on the center. Then, a line-mapping is performed using the automatized line-scanning tool. This tool allows the acquisition of a complete Raman spectrum at different points on a line by automatically moving the stage in one or two directions (x–y). The microscope objective used in this work, with a magniﬁcation of 20 , allows the visualization of a maximum area of 500 lm 70 lm (Fig. 1c). Therefore, in order to analyze the whole sample (10 mm) 20 lines with 5 equidistant points each have been measured. This was performed with the step-by-step procedure, in which the motorized stage is moved 500 lm (the line-mapping length) in the x direction to allow the analysis of the next part of the sample. The acquisition time for each spectrum was 100 s on an extended shift of 100–1200 cm 1. During the start of all Raman scan, a cosmic ray subtraction is automatically carried out to count any radioactive interference from the atmosphere or the sample. All spectra were re-calibrated daily with the emission lines of a Ne lamp. Spectra manipulation such as baseline adjustment,

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Fig. 1. (a) Macro-photograph of the studied sample, the square indicates the area of the analyzed sample. (b) Macro-photograph of the different eight regions in which we had divided the sample for further analysis. (c) Optical microscopy of the ﬁrst line-mapping.

smoothing, and normalization was performed using the Origin software. The overall appearance of the spectra obtained in this work was compared with the one of those in the RRUFF database [35]. 3. Results By using the Raman line-mapping procedure described before, we have obtained the following results, which can be divided into three parts: identification of the different secondary phases by using the Raman finger-print, analysis of the different regions and semi-quantitative analysis of the sample. The results obtained are coherent with the SEM–EDS study of the texture of the sample. 3.1. SEM–EDS analysis The texture of the uranium ore sample in polished section shows an alteration rim in contact with unaltered uraninite constituted by phase composed by U, O and C (Fig. 2b). The same secondary phase was found filling small fractures in the relicts of unaltered uraninite. The outer part of the altered uraninite is composed of a yellow secondary phase containing calcium, uranium and silicone (Fig. 2c) and irregular grains and inclusions containing lead, uranium and silicon (Fig. 2c). The microscopic observation of the texture suggest a first stage of uraninite alteration, characterized by the formation of oxide-carbonate phases and a second

phase This phase is characterized by the reaction with silica and calcium-rich fluids that affects the external zone of the sample, with formation of uranyl silicate phases. The exolution of radiogenic lead was observed by the formation of a separate lead bearing silicate phase. The Raman analysis of the sample is necessary to complete the fully characterization of the secondary mineral assemblage, as the EDS spectroscopy is unable to define the mineral phases observed, especially during the first stages of alteration of primary uraninite.

3.2. Identification of the secondary phases By scanning the sample with the Raman technique from the center outwards, four secondary uranium phases beyond uraninite were identified, by comparing the obtained spectra with the ones published by Frost: rutherfordine, UO2(CO3), soddyite, (UO2)2SiO4 2H2O, uranophane alpha Ca(UO2)2(SiO3OH)2 5H2O and kasolite, PbUO2SiO4 H2O, see spectra shown in Fig. 3. Rutherfordine was identified by means of the two symmetric stretching bands, m1, at 889 cm 1 and 1120 cm 1 of the (UO2)2+ and (CO3)2 groups, respectively. The band at 830 cm 1, is attributable to the m2 bending modes of the (CO3)2 group, m2(CO3)2 , and the band with lower intensity at 789 cm 1 is due to the m4 out of plane bending modes. In the low wavenumber region, we found three bands at 142 cm 1, 162 cm 1 and 220 cm 1 [36].

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Fig. 2. (a) Backscattered electron image of a section of uraninite sample showing replacement of uraninite by oxidized uranium phases. (b) EDS spectrum of the oxidecarbonate alteration product. (c) EDS spectrum of uranyl, calcium and silicate rich alteration zone that constitutes the main supergenic phase. (d) EDS spectrum of intergrown uranyl and lead silicate phase.

96 2 83

Soddyite

I(ν) / a.u.

Uranophane alpha

78 8 87 33

Rutherfordine

91 94 2 97 9 2

Kasolite

100

200

300

400

500

600

700

800

900

1000 1100 1200

ν / cm-1 Fig. 3. Raman spectra of the four secondary uranium phases identiﬁed: rutherfordine, UO2(CO3), uranophane alpha Ca(UO2)2(SiO3OH)2 5H2O, soddyite, (UO2)2SiO4 2H2O, and kasolite, PbUO2SiO4 H2O.

The uranyl silicate minerals found in the sample: uranophane alpha, Ca(UO2)2(SiO3OH)2 5H2O, soddyite, (UO2)2SiO4 2H2O, and kasolite PbUO2SiO4 H2O, were identiﬁed using the two internal modes m1 of the (UO2)2+ and (SiO4)4 . Symmetric stretching mode, m1(UO2)2+ at 798 cm 1 and m1(SiO4)4 at 967 cm 1 correspond to uranophane alpha and the expected overlapping of these bands, m1(UO2)2+ and m1(SiO4)4 at 832 cm 1 indicates soddyite. The Raman spectrum of kasolite has been characterized by the bands at 768 cm 1 and 912 cm 1, corresponding to the symmetric stretching modes m1(UO2)2+ and m1(SiO4)4 respectively, and the bands at 949 cm 1 and 972 cm 1 correspond to the m3(SiO4)4

bending modes. For the three uranyl silicates, the bands into the low wavenumber region, 200–300 cm 1 are assigned to m2 bending modes of (UO2)2+, whereas bands corresponding to the bending modes of (SiO4)4 , m2 and m4, appear at 400 cm 1 and 450– 600 cm 1, respectively [37,38]. The assignation to the different vibration modes of each phase are shown in Tables 1–4. The peak positions are in good agreement with the published literature values [36–39] and standard materials [35]. The ﬁngerprint used in this work to identify each phase in the sample was the symmetrical stretching vibration of the UO2+ 2 group, m1(UO2)2+. As it can be seen in the in Fig. 3, these bands are well resolved and do not overlap each other (frequencies indicated in bold in Tables 1–4). 3.3. Region analysis The distribution of the different phases along the sample can be divided into eight regions from the center outwards (see Fig. 1b), where a phase or a mixture of two or more phases predominates over the others.

Table 1 Rutherfordine UO2(CO3). Band

Assignation

a b c d

Not assigned Not assigned Not assigned m4(CO3)2

e f g

m2(CO3)2 m1(UO)2+ m1(CO3)2

Frequency This work 142 162 220 789 833 889 1120

Frequency [36] – – – 799 784 804 866 1115

300

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Table 2 Uranophane alpha Ca(UO2)2(SiO3OH)2 5H2O. Assignation

Frequency This work

Frequency [37]

a b c d

Not assigned m2(UO2)2+ m2(UO2)2+ m2(UO2)2+

170 209 254 300

e f g h i

m2(SiO4)4 m4(SiO4)4 m4(SiO4)4 m1(UO2)2+ m1(SiO)4

403 473 547 798 967

166.7 213.7 250.3 306.5 288.9 280.5 398.9 469.5 544.6 796.9 963.9

0.8 mm

soddyite

400

600

800

1000

1200

ν / cm-1 Fig. 4. Raman spectra obtained in region 2 and 3 at 0.8 and 0.9 mm from the center, of the sample respectively which show soddyite for region 2 and a mixture of soddyite and rutherfordine in region 3.

rutherfordine + soddyite + uranophane alpha +kasolite

1.7mm

3.3.2. Region 4 Region 4 (1.0–3.3 mm) is characterized by the coexistence of the four secondary phases, soddyite, rutherfordine, uranophane alpha and kasolite in different proportions, (see Fig. 5). The proportions of the different phases in the mixtures can be compared by the analysis of the different relative intensities of the spectra bands. As it is well known, Raman spectroscopy can be used as an analytical technique to extract quantitative information [40]. The intensity of Raman scattering, IR, can be written as

1.6mm

1.3mm 200

ð1Þ

where IL is the laser intensity, r is the Raman cross-section or scattering efﬁciency, g is an instrument parameter, P is the sample path length, and C is the concentration [41]. Therefore, intensity peak ratios (IA/IB) may be used to determine relative concentrations RCA/CB of two components, A and B; thus, RCA/CB = CA/CB, where CA and CB are the concentrations of A and B, respectively. Due to the fact that the cross sections of the different compounds are not the same, it is not possible to calculate the concentration of each component in the mixture, but it is possible to compare the concentration of each component in different mixtures. Then, by using the different peak ratios one can conclude that the spectra acquired at 1.7 and 1.3 mm shown in Fig. 5 correspond to mixtures of these four phases where the amount of soddyite is higher than in the mixture corresponding to spectra acquired at 1.6 mm; i.e. the relative intensity of

200

Intensity / a.u.

3.3.1. Region 1, 2 and 3 The region 1 extends approximately from the center of the sample (0 mm) to 0.4 mm and is considered the core of the sample, comprising uraninite (UO2+x), without any alteration products. The region 2 (0.4–0.8 mm) is characterized by the presence of soddyite and the region 3 (0.8–1.0 mm) corresponds to a mixture of soddyite and rutherfordine (spectra at 0.8 and 0.9). Fig. 4 shows the Raman spectra obtained in region 2 and 3 at 0.8 and 0.9 mm from the center, of the sample respectively. Soddyite has been identiﬁed in the region 2 and a mixture of soddyite and rutherfordine in region 3.

IR ¼ ðIL rgPÞC;

0.9 mm (region 3)

soddyite + rutherfordine

Intensity / a.u.

Band

400

600

800

1000

1200

ν / cm-1 Fig. 5. Raman spectra obtained in region 4 at 1.3, 1.6 and 1.7 mm from the center of the sample, which show a mixture of rutherfordine, soddyite, uranophane alpha and kasolite.

the band m1(UO2)2+ of soddyite, Im1(soddyite), in relation to the band intensities of the m1(UO2)2+ of the others minerals, Im1(rutherfordine), Im1(uranophane alpha), and Im1(kasolite), is higher in the spectra at 1.7 and 1.3 mm than in the spectra corresponding to 1.6 mm. 3.3.3. Region 5 and 6 Fig. 6 shows the typical spectra found in region 5 (3.3–4.6 mm) and 6 (4.6–7.1 mm). In these two regions, from 3.3 mm to the center of the sample, the uranyl carbonate, identiﬁed as rutherfordine,

Table 4 Kasolite PbUO2SiO4 H2O.

Table 3 Soddyite (UO2)2SiO4 2H2O. Band

Assignation

Frequency This work

Frequency [38]

Band

Assignation

Frequency This work

Frequency [31]

a b c d

Not assigned m2(UO2)2+ m2(UO2)2+ m2(UO2)2+ m2(UO2)2+ m2(SiO4)4 m4(SiO4)4 m1(UO2)2+ + m1(SiO4)4

107 195 225 293 312 404 463 832

111 190 229 290 310 – 459 828

a b c d

Not assigned

m2(UO2)2+ m2(UO2)2+ m2(SiO4)4

107 217 237 424

e f g

m2(SiO4)4 m1(UO2)2+ m1(SiO4)4

553 768 912

107.5 217.7 234.3 454.6 415.1 550.4 766.7 903.6

e f

301

L.J. Bonales et al. / Journal of Nuclear Materials 462 (2015) 296–303

soddyite + uranophane alpha

uranophane alpha

8.6 mm

3.7 mm

Intensity / a.u.

3.8 mm

8.5 mm

3.6 mm

200

400

600

800

1000

8.4 mm

200

1200

400

600

ν/

ν / cm-1

soddyite + uranophane alpha + kasolite

800

1000

uranophane alpha + kasolite

9.3 mm Intensity / a.u.

Intensity / a.u.

4.8 mm

4.7 mm

1200

cm-1

9.2 mm

4.2 mm

9.1 μm 200

400

600

ν/

800

1000

1200

cm-1

Fig. 6. (a) Raman spectra obtained in region 5 at 3.6, 3.7 and 3.8 mm from the center of the sample which show the characteristic bands of the soddyite and uranophane alpha mixture. (b) Raman spectra obtained in region 6 at 4.2, 4.7 and 4.8 mm from the center of the sample which show a mixture of soddyite, uranophane alpha and kasolite.

was completely absent and substituted by uranyl silicates, such as soddyite and uranophane alpha in region 5. These silicates, have been identiﬁed as kasolite in the region 6. The analysis of the relative intensities in these regions indicates different proportions of the different phases in the sample. Fig. 6a shows the mixture corresponding to the spectrum at 3.8 mm, it has a relation of soddyite/uranophane alpha higher than the mixtures of the spectra at 3.6 and 3.7 mm. Moreover the mixture corresponding to the spectrum at 4.7 mm shown in Fig. 6b is a mixture richer in kasolite than in the others mixtures. 3.3.4. Region 7 and 8 In Fig. 7 the spectra analyzed in region 7 (7.1–8.8 mm) and 8 (8.8–1.0 mm) are shown. In these two outer regions the predominant phase is uranophane alpha, being this the only phase present in region 7, while in region 8 this phase coexists with kasolite at different proportions. As it can be seen in Figs. 4–7, most of the Raman spectra acquired in this work correspond to mixtures of different phases in which the fraction of the different minerals is highly variable, very typical of gummites as expected. 4. Discussion In order to perform a semi-quantitative analysis of the sample in order to identify the presence or absence of different phases

200

400

600

800

1000

1200

ν / cm-1 Fig. 7. (a) Raman spectra obtained in region 7 at 8.4, 8.5 and 8.6 mm from the center of the sample which show the characteristic spectra of uranophane alpha. (b) Raman spectra obtained in region 8 at 9.1, 9.2 and 9.3 mm from the center of the sample which show a mixture of uranophane alpha and soddyite.

along the different regions of gummite (from 0.4 to 10.00 mm), 100 spectra have been processed as will be explain below. It should be mentioned that the quantitative methods had been developed initially for gases, and then gases dissolved in fluid inclusions and have been adapted here for use in solid mineral mixtures. The performed analysis is based on a characteristic of Raman spectra for mixtures: the spectra can be understood as the direct sum of the individual spectrum of each component in the mixture as long as these components do not interact with each other. Therefore, the vibration bands do not undergo any displacement, and the band profile of the mixture spectra results in the spectra of the different components or vice versa. In order to calculate the number of contributions of a given band, which is not always possible to the naked eye, it has been analyzed the resulting spectrum by the second derivative method [42]. The first derivative gives us an idea of the number of contributions involved but, as usual in spectroscopy, is the second derivative which enables us to determine the number of contributions, since each one leads to a minimum. As an example, in Fig. 8, the analysis of the second derivative of the m1(UO2)2+ stretch region at 700–900 cm 1 is shown. This figure highlights that when the amount of a phase in a mixture is very small, in proportion to the other present phases, it is necessary to perform the analysis of the second derivative to identify the number of contributions, (Fig. 8), since the band corresponding to a lower amount appears as a shoulder, and not as a resolved peak. Thereby, we

L.J. Bonales et al. / Journal of Nuclear Materials 462 (2015) 296–303

Intensity / a.u.

second derivative of Intensity / a.u.

302

700

720

740

760

780

ν/

800

820

840

860

700

720

740

cm-1

760

780

800

820

Raman shift /

840

860

880

cm-1

Fig. 8. (Left) Open points show the Raman spectra corresponding of a mixture of two minerals, lines represent the best ﬁtting to a two Gaussian curves. (Right) The second derivate of the Raman spectra.

determine the number of contributions by calculating the second derivate of each spectrum and by constructing a data matrix of 0 and 1, where 0 means there is no minimum to the characteristic frequency of the mineral, and 1 means there is a minimum at the characteristic frequency of the mineral. Fig. 9 shows the diagrams constructed by this method, i.e. we plot the presence (1) or absence (0) for each phase vs. the position of the analyzed point, from the center of the sample outwards (0–10 mm). Lines are the smoothed data and indicate the trends of increase or decrease of each phase along the sample. As it can be seen in Fig. 9, the center of the sample, 0–0.4 mm, is composed by uraninite. The rutherfordine is the predominant phase in the inner part, 0.4–3.3 mm, in contact with the uraninite core, and then is absent from 3.3 mm. The analysis of the next region indicates a mixture of uranyl silicates: soddyite, uranophane alpha and kasolite. Soddyite prevails in the inner part, 0.4–7.1 mm; uranophane alpha predominates in the outer part of the sample, 7.1–10 mm, and kasolite appears intermittently (1.0–3.3 mm; 4.6–7.1 mm and 8.8–10 mm). It should be noted that schoepite, (UO2)4O(OH)6 6H2O, the expected uranyl phase formed by corrosion of uraninite under atmospheric conditions [15] or by silica-poor meteoric waters, has not been observed in the sample analyzed in this work. The absence of a signiﬁcant occurrence of schoepite in the sample could be explained by its rapid transformation, to rutherfordine,

Soddyite

Uranphane alpha

Kasolite Rutherfordine

X (mm) Fig. 9. Points indicate the presence (1) or absence (0) for each analyzed mineral vs. position of the analyzed point. Lines indicate the trends of increase or decrease of each phase along the sample.

which appears as replacement structures in the gummite rim, suggesting that the original schoepite or metaschoepite has been replaced by the rutherfordine, as it is the stable phase in CO2 rich fluids in subsurface conditions [43]. The formation of rutherfordine, confined in the inner zone of the corrosion rim of uraninite, could be one of the first steps of alteration, after the formation of the schoepite or metaschoepite or other oxy-hydroxydes with different U(IV)–U(IV) proportions. The next alteration products are the uranyl silicates soddyite and uranophane alpha. Soddyite is the first silicate precipitated by reaction of silicate rich solutions with uraninite and the first alteration product, replacing them in the vicinity of the primary mineral. The formation of soddyite or uranophane depends on the activity ratio (Ca)/(H+). As a result, Ca poor and low pH waters favor the replacement of schoepite by soddyite [43]. The formation of uranophane alpha requires a calcium and silica rich fluid, provided by the alteration of the feldspars (mainly plagioclases) that usually surround the uraninite crystals in the pegmatite [34]. Hence, the circulant calcium and silica rich water determines the distribution of silicates in the gummite. The external zone of the corrosion rim, in contact with altered feldspars, is almost entirely composed by uranophane and the inner zone, in contact with uraninite, is more U rich and Ca poor, and is dominated by soddyite. The mineral assemblage, in presence of persisting uraninite, is determined by the composition of infiltrating waters. An interesting feature of our sample is the lead enrichment in the form of kasolite in the gummite zone. The accumulation of lead in the uraninite destabilizes its structure by induction of strain. Under oxidizing conditions, lead combines with uranyl to form Pb-uranyl minerals. Lead is not incorporated to rutherfordine, soddyite and uranophane but instead kasolite accumulates in fractures and veinlets. The role of radiogenic Pb is essential in the formation of secondary phases. The formation of vandendriesscheite, coetaneous with the formation of schoepite, alters incongruently in presence of CO2 waters to form uranyl carbonates and lead enriched phases, as masuyite, whose alteration in silica rich waters leads to the formation of kasolite. The lower mobility of Pb mineral phases compared with uranyl phases leads to a gradual enrichment in kasolite on the gummite. 5. Conclusion In this work we present the Raman spectra of the alteration products of a uraninite sample, (an analogue of the spent fuel), taken from the Sierra Albarrana, Spain. The identification of the

L.J. Bonales et al. / Journal of Nuclear Materials 462 (2015) 296–303

different secondary phases, have been performed by the analysis of the symmetrical stretching vibration of the uranyl group (UO2+ 2 ), taken as ﬁngerprint of the found phases: rutherfordine, UO2(CO3), soddyite, (UO2)2SiO4 2H2O, uranophane alpha Ca(UO2)2(SiO3OH)2 5H2O and kasolite, PbUO2SiO4 H2O. The spatial and temporal sequence of alteration products obtained was: (1) uraninite constitutes the unaltered core of the sample, 0–0.4 mm. (2) Rutherfordine appears in the inner part, 0.4–3.3 mm, in contact with the uraninite core. (3) Then a mixture of uranyl silicates, soddyite, uranophane alpha and kasolite are found. Soddyite prevails in the inner part, 0.4–7.1 mm; uranophane alpha predominates in the outer part of the sample, 7.1–10 mm, and kasolite appears intermittently (1.0–3.3 mm; 4.6–7.1 mm and 8.8–10 mm). This sequence had been obtained by using a semi-quantitative analysis developed in this work, which enables to elucidate the presence or absence of the different phases in an easy and quick way and moreover, without using other complementary techniques. Schoepite, (UO2)4O(OH)6 6H2O, the expected uranyl phase formed by corrosion of uraninite under atmospheric conditions, has not been observed due to its rapid transformation to rutherfordine, which appears as replacement structures in the gummite rim. Because the knowledge of Raman spectra of uranyl-based minerals is still rather limited, this study, as a part of our ongoing research into the use of Raman spectroscopy, intends to increase the Raman database spectra of uranium based-minerals, as important in the ﬁeld of nuclear waste disposal. Acknowledgments Authors thank to Jose González del Tánago for many useful comments and to the Museo de Ciencias Naturales de Alava for kindly providing the samples used in this study. This work was supported by ENRESA in the Project: No. 079000189 entitle ‘‘Aplicación de técnicas de caracterización en el estudio de la estabilidad del combustible nuclear irradiado en condiciones de almacenamiento’’ (ACESCO). References [1] SKB 91, Final Disposal of Spent Nuclear Fuel. Importance of the Bedrock for Safety, SKB Report 92-20, May 1992. [2] D.W. Shoesmith, J. Nucl. Mater. 282 (2000) 1–31. [3] S. Sunder, Nucl. Technol. 122 (1998) 211–221. [4] S. Sunder, Alpha, Beta and Gamma Dose Rates in Water in Contact with Used CANDU UO2 Fuel, Atomic Energy of Canada Ltd., Report, AECl-11380, COG-95340, 1995. [5] W. Gray, Effect of Surface Oxidation, Alpha Radiolysis and Salt Brine Composition on Spent Fuel and UO2 Leaching Performance, PNL/SRP-6689, 1988, pp. 4.6–4.8.

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[6] R. Wang, J.B. Katayama, Nucl. Chem. Wast. Manage. 3 (1982) 83–90. [7] J. Bruno, E. Cera, L. Duro, T.E. Eriksen, L.O. Werme, J. Nucl. Mater. 238 (1996) 110–120. [8] J. de Pablo, I. Casas, J. Giménez, V. Marti, M.E. Torrero, J. Nucl. Mater. 232 (1996) 138–145. [9] I. Casas, J. Giménez, V. Martí, M.E. Torrero, J. de Pablo, Radiochim. Acta 66–67 (1994) 23–27. [10] P. Fors, P. Carbol, S. Van Winckel, K. Spahiu, J. Nucl. Mater. 394 (2009) 1–8. [11] J. Bruno, I. Casas, M.C.A. Sandino, J. Nucl. Mater. 190 (1992) 61–69. [12] K. Ollila, J. Nucl. Mater. 190 (1992) 70–77. [13] D.E. Grandstaff, Econ. Geol. 71 (1976) 1493–1506. [14] J. Janeczek, R.C. Ewing, J. Nucl. Mater. 190 (1992) 157–173. [15] R.J. Finch, R.C. Ewing, J. Nucl. Mater. 190 (1992) 133–156. [16] J. Janeczek, R.C. Ewing, J. Nucl. Mater. 185 (1991) 66–77. [17] J. Janeczek, R.C. Ewing Janeczek, J. Nucl. Mater. 190 (1992) 128–132. [18] J. Janeczek, R.C. Ewing, V.M. Oversby, L.O. Werme, J. Nucl. Mater. 238 (1996) 121–130. [19] E.C. Pearcy, J.D. Prikryl, W.M. Murphy, B.W. Leslie, Appl. Geochem. 9 (1994) 713–732. [20] D. Zhao, R.C. Ewing, Radiochim. Acta 88 (2000) 739–749. [21] A.P. Deditius, S. Utsunomiya, R.C. Ewing, Geochim. Cosmochim. Acta 71 (2007) 4954–4973. [22] L. Perez del Villar, J. Bruno, R. Campos, P. Gomez, J.S. Cozar, A. Garralon, B. Buil, D. Arcos, G. Carretero, J. Ruiz Sanchez-Porro, P. Hernan, Chem. Geol. 190 (2002) 395–415. [23] C. Frondel, US Geol. Surv. Bull. 1064 (1958). [24] C. Frondel, Am. Mineral. 41 (1956) 539–568. [25] J. Dubessy, M.-C. Camon, F. Rull, Raman Spectroscopy applied to Earth Sciences and Cultural Heritage. Editors. The Mineral Society of Great Britain & Ireland, London, 2012. [26] C.C. Allen, L.S. Butler, N. Anh Tuan, J. Nucl. Mater. 144 (1987) 17–19. [27] B.S.M. RaO, E. Ganter, J. Reinhart, D. Steinert, H.J. Ache, J. Nucl. Mater. 170 (1990) 39–49. [28] M. Amme, R. Renker, B. Schimjid, M.P. Feth, H. Bertagnolli, W. Döbelin, J. Nucl. Mater. 306 (2002) 202–212. [29] J. González del Tánago, M. Martinez, M. Peinado, I Congreso Español de Geología II (1984) 131–145. [30] F. Tornos, C.M.C. Inverno, C. Casquet, A. Mateus, G. Ortiz, V. Oliveira, J. Iberian Geol. 30 (2004) 143–181. [31] J. González del Tánago, Boletín Sociedad Española de Mineralogía 14–1 (1991) 54–55. [32] J. González del Tánago, M. Peinado, J.L. Brändle, Boletín Sociedad Española de Mineralogía 14–1 (1991) 105–106. [33] P. Cherny, T.S. Ercit, Can. Mineral. 43 (2005) 2005–2026. [34] J. González del Tánago, PhD. Thesis. Universidad Complutense de Madrid, 1993. [35] R.T. Downs. The RRUFF Project: an integrated study of the chemistry, crystallography, Raman and infrared spectroscopy of minerals. Program and Abstracts of the 19th General Meeting of the International Mineralogical Association in Kobe, Japan, 2006. O03-13. [36] R.L. Frost, C. Jiri, J. Raman Spectrosc. 40 (2009) 1096–1103. [37] R.L. Frost, J. Cejka, M.L. Weier, W. Martens, J. Raman Spectrosc. 37 (2006) 538– 551. [38] R.L. Frost, M.L. Weier, W. Martens, T. Kloprogge, J. Cejka, Spectrochim. Acta 63 (2006) 305–312. [39] B.M. Biwer, W.L. Ebert, J.K. Bates, J. Nucl. Mater. 175 (1990) 188–193. [40] J.D. Pasteris, B. Wopenka, J.C. Seitz, Geochim. Cosmochim. Acta 52 (1988) 979– 988. [41] S. White, Appl. Spectrosc. 67 (2010) 819–827. [42] E. del Corro García, pHD Thesis. Universidad Complutense de Madrid. CC. Químicas Departamento de Química Física I, 2011. [43] R. Finch, T. Murakami, Rev. Mineral. 38 (1999) 91–179.

ARTICLE Received 2 Sep 2013 | Accepted 7 Jul 2014 | Published 6 Aug 2014

DOI: 10.1038/ncomms5600

Formation of recent Pb-Ag-Au mineralization by potential sub-surface microbial activity Fernando Tornos1, Francisco Velasco2, Ce´sar Menor-Salva´n1, Antonio Delgado3, John F. Slack4 & Juan Manuel Escobar5

Las Cruces is a base-metal deposit in the Iberian Pyrite Belt, one of the world’s best-known ore provinces. Here we report the occurrence of major Pb-Ag-Au mineralization resulting from recent sub-surface replacement of supergene oxyhydroxides by carbonate and sulphide minerals. This is probably the largest documented occurrence of recent microbial activity producing an ore assemblage previously unknown in supergene mineralizing environments. The presence of microbial features in the sulphides suggests that these may be the ﬁrst-described natural bacteriomorphs of galena. The low d13C values of the carbonate minerals indicate formation by deep anaerobic microbial processes. Sulphur isotope values of sulphides are interpreted here as reﬂecting microbial reduction in a system impoverished in sulphate. We suggest that biogenic activity has produced around 3.1 109 moles of reduced sulphur and 1010 moles of CO2, promoting the formation of ca. 1.19 Mt of carbonates, 114,000 t of galena, 638 t of silver sulphides and 6.5 t of gold.

1 Centro de Astrobiologıá. Ctra Ajalvir km. 4.5, Torrejon de Ardoz, 28850 Madrid, Spain. 2 Dpto. Mineralogıá y Petrologıá, Facultad de Ciencia y Tecnologıá, Universidad del Paı´s Vasco UPV/EHU, 48080 Bilbao, Spain. 3 Laboratorio de Biogeoquı´mica de Iso´topos Estables, Instituto Andaluz de Ciencias de la Tierra IACT (CSIC-UGR). Avda. de las Palmeras, 4, 18100 Armilla, Granada, Spain. 4 U.S. Geological Survey, National Center, MS 954, Reston, Virginia 20192, USA. 5 Cobre Las Cruces S.A., Gerena, 41860 Seville, Spain. Correspondence and requests for materials should be addressed to F.T. (email: f.tornos@csic.es).

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ne of the principal ecological niches for extremophilic microbes is the sub-surface environment, where there is more opportunity for the existence of anaerobic life1. Here, the absence of oxygen is a limiting factor for heterotrophic processes. Consequently, lithoautotrophic bacteria create microdomains with metastable mineral assemblages that are inhibited under low-temperature, oxygenated surface conditions. Recent studies suggest that nearly the same amount of biomass is present in subterranean and sub-seaﬂoor settings as on Earth’s surface1–4. The contact between the Late Devonian–Early Carboniferous (ca. 360–330 Ma) volcanogenic massive sulphide deposit in the Las Cruces mine (Seville, southern Spain) and unconformably overlying Cenozoic sedimentary rocks of the Guadalquivir basin (Fig. 1a) hosts an unusual secondary mineral assemblage that includes galena, carbonates, iron sulphides, silver-rich sulphides and gold5–8. This assemblage forms a sub-horizontal rock layer, B3–10 m thick, which caps the tilted primary massive sulphide lens. Beneath the layer is a large secondary cementation zone, up to 40 m thick, containing abundant chalcocite that constitutes the ore currently mined at Las Cruces, making it one of the richest copper deposits worldwide (17 Mt at 6.9% Cu)7. At the mine, the basement-cover contact is currently the locus of a major stratiform aquifer dominated by sulphate-bearing, calciumbicarbonate, neutral to alkaline groundwaters. The massive sulphide orebody is similar to those elsewhere in the southern part of the Iberian Pyrite Belt9,10 that formed in the latest Famennian (ca. 360 Ma) and subsequently were folded and metamorphosed during the Variscan orogeny (330–300 Ma). The Las Cruces deposit differs, however, in occurring below 150 m of Late Tortonian–Messinian (ca. 7.2–5.3 Ma) sandstone and marl. This setting hindered its discovery and exploitation until 1994, but contributed greatly to the formation of the unusual mineral assemblages.

Other massive sulphide deposits of the Iberian Pyrite Belt that were exhumed in Miocene time (o23 Ma) underwent extensive sub-aerial supergene alteration and have classical secondary alteration zones11. The available data suggest that some of these secondary zones formed at 7–8 Ma (ref. 11). Here, sulphides were weathered and the leached caps were depleted of Cu and Zn but enriched in Fe, producing the well-known orange and red gossans that overlie most of the exposed volcanogenic massive sulphide deposits of the Pyrite Belt, and that contain abundant and visually distinctive haematite and goethite. Within these deposits, the contact between the cementation zone and the gossan is the locus of a yellowish layer, several metres thick, enriched in Pb, As, Sb, Hg, Ag and Au, and having significant amounts of barite and quartz. The lead is found in cerussite, anglesite, plumbojarosite or beudantite, whereas the silver seems to occur as small amounts of argentojarosite; gold is present as sub-microscopic inclusions within the goethite11. At Las Cruces, this typical gossan has been replaced by a younger sulphide-carbonate assemblage, which is, to our knowledge, foreign to sub-aerial supergene environments. We interpret this unusual rock as having formed due to subsurface metabolism of anaerobic prokaryotes under the marl unit that seals the deposit. Results The secondary zone. In great contrast to the general behaviour of sub-aerially exposed orebodies in the Iberian Pyrite Belt and elsewhere, the gossan at Las Cruces shows a complex sequence of post-formational modifications (Fig. 1). Here, the original goethite- and haematite-rich assemblage is replaced by a siderite- and galena-rich rock (Red Rock), which in turn is replaced by a younger, highly heterogeneous rock (Black Rock) having variable proportions of coarse-grained (0.5–2 mm) calcite, fine-grained (o10 mm) galena and minute acicular to globular aggregates Las Cruces secondary profile

b Typical IPB gossan profile

PUL O DO L OB O GROUP

Iberian Pyrite Belt

Las Cruces Seville

Huelva H UELVA

PORTUGAL SPAIN Faro

Baixo Alentejo Flysch Group Volcanic Sedimentary Complex Phyllite-Quartzite Group

2 .Gossan (mainly Miocene) Iberian Pyrite Belt

3. Jarosite level (lower gossan) 4. Cementation zone 5. Massive sulphides

Cementation system

1 .Over burden (Holocene)

Massive sulphide deposits Cropping out deposits Deep deposits

m (approx)

Post-Variscan cover Granitic intrusives

50 km

Late Devonian

280.000

3 0

Marl

Sandstone

2 Biogenic system

Miocene

360.000

Aracena

Messinian

BBeja EJ A

320.000

340.000

Miocene-present day?

sin

300.000

Late Devonian

260.000

220.000

Red Rock (Pb) Black Rock (Pb-Ag-Au) Barren pyrite

Cementation zone (Cu)

Massive sulphides (Cu-Zn-Pb)

Figure 1 | Geologic setting of the Las Cruces mine. (a) Regional geological map showing the location of the deposit. Adapted from ref. 10 with permission from Elsevier; (b) idealized sections of the typical gossan of the Iberian Pyrite Belt and of the secondary mineralized zone at Las Cruces (graphical scale, ca. 20 m). Outcropping ore deposits of the Iberian Pyrite Belt have well-formed gossans. Las Cruces is an exception, with a secondary zone covered by Cenozoic sediments. The quoted ages are depositional ages for the primary massive sulphides and the sediments but are formational ages of the gossan and secondary mineralization. 2

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composed of greigite and smythite. The Black Rock also contains minor amounts of intergrown acanthite, sternbergite, proustitexanthoconite, pearceite, jamesonite, cinnabar and cassiterite; gold-rich alloys are intergrown with the carbonates and galena. Barite and quartz are common gangue minerals in both the Black and Red rocks. Thin (o1 mm) sub-vertical veins of galena penetrate through the overlying marl12 and also cut the underlying massive sulphides. Blake8 ﬁrst suggested that both the Black and Red rocks were of bacterial origin. Bacteriomorphic structures. The Black Rock contains abundant, unbranched and twisted, thread-like aggregates of galena ca. 5–10 mm in length and r1 mm radius located on carbonate crystals (Fig. 2). These aggregates are morphologically similar to the microbial structures reported previously in the laboratory13 and in fossil14,15 geomicrobiological systems. The lack of external morphological features suggests that the galena precipitated as an extracellular polymer coating of bacteria as a result of metabolic sulphate reduction16. No previous records exist in natural systems of bacteriogenic structures composed of galena. Biogenic galena, to our knowledge, has only been described as a product of the bacterially mediated reduction of anglesite in discarded batteries17. Stable isotope geochemistry. Carbon, oxygen and sulphur isotope analyses were obtained on the Red and Black rocks, as well

Figure 2 | Bacteriomorphic galena on calcite in the Black Rock. (a) Scanning electron microscope (SEM) secondary electron image (scale bar, 10 mm); (b) SEM backscattered electron image (scale bar, 5 mm).

as on the overlying marl and local groundwaters (Table 1). The d13C and d18O values of the siderite and calcite range from 42 to 18% versus V-PDB (d13C) and þ 22.1 to þ 27.5% versus V-SMOW (d18O). The d34S values of the sulphides (versus V-CDT) are þ 11.9 to þ 25.9% for galena and þ 16.3 to þ 19.5% for greigite–smythite. Discussion The strict localization of the Red and Black rocks in the contact zone between the paleosurface and the deformed massive sulphides, and the presence of veins of galena cutting the overlying Tertiary sediments, indicate that these rocks are not related to the in situ submarine oxidation of the massive sulphides shortly after their formation at or near the seafloor, or by much later sub-aerial processes as proposed by Knight5. More probably, the Red and Black rocks formed after burial beneath the Late Tortonian–Messinian sediments8. Both the presence of bacteria replaced by galena and the stable isotope data are consistent with replacement of the gossan being related to biogenic processes. The d13Ccarbonate values are best interpreted as reflecting the mixing between 13C-depleted carbon and relatively isotopically heavy carbon (d13C4 9%), with this latter end member derived from dissolved inorganic carbon (DIC; defined as SC (CO2,aq þ H2CO3 þ HCO3 )) in local groundwaters (Fig. 3a and Table 1). The most plausible source of the isotopically light carbon is the in situ biogenic oxidation of methane and other light hydrocarbons (Fig. 4); this interpretation is supported by the accumulations of light gas that are common in the Guadalquivir basin18. The ultimate origin of the light carbon may be thermal maturation of Palaeozoic shale within the underlying basement, a process that released gas that then ascended through fractures, accumulating below the sealing marl unit. Interpretation of the d34S values for the sulphides is not straightforward, because the measured d34S values are strikingly high compared with d34Ssulphate values of the groundwaters currently flowing into the open pit of the Las Cruces mine ( þ 6.3 to þ 15.2%) and with those of the underlying massive sulphides ( 6.8 to þ 8.2%)8,19 (Fig. 3b and Table 1). However, these high d34S values are similar to those of associated barite ( þ 20.4%)5. This situation is rather unusual for biogenically mediated, sulphide-rich ore systems20, because strongly negative d34S values are generally regarded as prima facie evidence of biogenic sulphate reduction21,22, whereas small DSO4-H2S values are characteristic of thermochemical sulphate reduction23. At the low temperatures expected in the secondary environment of the Las Cruces deposit (Eo70–80 °C), sulphate reduction can only be mediated biogenically, typically by dissimilatory anaerobic processes, whereas abiotic sulphate reduction is kinetically inhibited23. Some bacteria are known to excrete reduced sulphur with small DSO4-H2S values24. However, we interpret the isotopically high d34S values as recording fractionation in a system having excess organic matter (TOC4 4SO24 ) and high rates of reduction, whereby all of the sulphate was rapidly transformed into sulphide24,25. The isotope values reflect the complete consumption of sulphate during anaerobic oxidation of electron donors in a closed system26, a process that is consistent with the high d34S values of the sulphides. The presence of iron monosulphides is indicative of a low-fS2 system and a limited availability of reduced sulphur. The ultimate origin of the groundwater sulphate is unknown, but it could have a seawater origin27, derived either from recent seawater intruding deep aquifers in the basin and upwelling along faults, or from the dissolution of accessory sulphate interbedded with the marl. The mineral assemblages of siderite–calcite and sulphides are unstable under sub-aerial oxic conditions, and in the Las Cruces

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Table 1 | Stable isotope composition of groundwater and carbonates and sulphides from Las Cruces. Identiﬁer EXT105 EXT201 EXT206 EXT305 EXT308 EXT502A EXT505 EXT603 LCW 2 LCW 3 LCW 5 LCW 6 LCW 7 LCM-182 LCM-183 LCM-184 LCM-171 LCM-013 LCM-035 LCM-165 LCM-166 * * * 67 69 89 LCM-023 LCM-028 LCM-039 LCM-040 LCM-042 LCM-063 LCM-229 LCM-229 LCM-417 LCM-418 LCM-419 LCM-420 LCM-421 LCM-422 LCM-423 LCM-424 LCM-425 LCM-426 LCT-15.1 LCT-15.2 LCT-16.1 LCT-16.2 LCT-20.1 LCT-20.2 LCT-21.1 CLC-43 LCT-771 LCT-772

Description Open pit-aquifer Open pit-aquifer Open pit-aquifer Open pit-aquifer Open pit-aquifer Open pit-aquifer Open pit-aquifer Open pit-aquifer Open pit-aquifer Open pit-aquifer Open pit-aquifer Open pit-aquifer Open pit-aquifer Tertiary marl Tertiary marl Tertiary marl Tertiary calcarenite Red Rock Red Rock Red Rock Red Rock Red Rock Red Rock Red Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock Black Rock

Sample Water Water Water Water Water Water Water Water Water Water Water Water Water Calcite Calcite Calcite Calcite Siderite, galena Siderite, galena Siderite, galena Siderite, galena Siderite Siderite Siderite Galena Galena Galena þ greigite Galena Galena Galena Galena Calcite, galena Galena Calcite Galena Calcite, galena þ greigite Calcite, galena Calcite, galena Calcite, galena Calcite, greigite Calcite, galena Calcite, galena Calcite, galena Calcite, galena Calcite, galena Greigite Greigite Greigite Greigite Greigite Greigite Greigite Greigite Galena Galena

d18O (% VSMOW) 5.1 5.2 5.0 5.1 5.2 5.3 5.3 5.5 5.8 5.8 5.5 4.5 4.2 32.3 33.6 30.8 32.8 25.1 22.8 23.8 24.0 27.5 26.0 25.8

d13C (% VPDB) 7.9 7.7 8.8 6.3 8.8 9.0 8.7 8.6 7.6 8.6 8.3 7.0 6.8 1.8 1.3 0.9 5.3 27.9 22.4 17.6 18.5 35.5 33.0 41.7

22.1

23.8

22.2

30.9

22.1 23.8 23.4 23.3 24.9 22.2 22.9 24.6 24.8 23.0

36.9 21.1 26.7 20.1 23.2 37.3 29.7 31.7 26.7 32.4

dD (% VSMOW) 31.3 31.2 30.7 30.4 32.3 32.0 33.0 34.5 29.2 31.3 30.8 30.4 31.1

d34S (% VCDT) 6.3 6.5 7.9 15.1 15.2

22.0 21.7 12.5 19.1

21.0 19.9 17.4 20.4 19.8 24.4 25.9 21.0 17.5 19.0 21.5 22.7 22.0 22.9 19.2 18.6 19.1 17.8 16.3 11.9 16.8 17.8 16.7 16.3 19.1 19.5 17.9 16.9 13.2 12.3

VCDT, Vienna Canyon Diablo Troilite; VPDB, Vienna-PDB; VSMOW, Vienna Standard Mean Oceanic Water. *Samples from Blake8.

deposit could only have formed in an aqueous environment enriched in CO2 and, to some extent, in H2S. Thus, ﬂuid–rock interactions should be able to produce three superimposed processes to form the observed assemblages: an increase of fCO2 with stabilization of carbonates; a redox change with reduction of Fe3 þ to Fe2 þ , also related to the replacement of goethite/ haematite by siderite; and sulphidization accompanied by the deposition of the sulphides. 4

The evolution of the gossan to the Red Rock must involve a complex set of reactions summarized as follows: CH4;aq þ SO24 þ H þ , HCO3 þ H2 Saq þ H2 O 8 goethiteðFeOOHÞ þ CH4;aq þ 15H , 8 Fe2 þ þ HCO3 þ 13 H2 O HCO3

2þ

þ Fe

ð1Þ

, sideriteðFeCO3 Þ þ H

ð2Þ þ

ð3Þ

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δ18O (‰ VSMOW)

5 0

δ13 C (‰ VPDB)

–5 –10 –15

presence of organic compounds capable of reducing Fe3 þ , the extent of abiotic iron reduction is much less than that produced by biological reduction28 and, moreover, the actual role of phenolic compounds in direct abiotic reduction, or as electron shuttles from bacteria to Fe3 þ mineral surfaces, is controversial. The H2S generated during reaction (1) should increase fS2, leading to the destabilization of the cerussite and precipitation of galena H2 Saq þ cerussiteðPbCO3 Þ , galenaðPbSÞ þ CO2;aq þ H2 O

–20 Tertiary marl and sandstone (cc)

–25

ð4Þ

Calcite equil NP aquifer –30

Red Rock (sid)

–35

Black Rock (cc) Massive sulphides (sid-ank)

–40

Finally, the formation of the Black Rock is related to replacement of the siderite by calcite: sideriteðFeCO3 Þ þ H2 Saq þ Ca2 þ , iron monosulphideðFeSÞ þ 2H þ þ CaCO3

18 16

Massive sulphides

In a strict sense, reactions (1) and (2) are not needed for the stabilization of the carbonate, because the ﬂowing groundwater carries CO2 and, hence, reaction (1) could be replaced by

Black Rock

Red Rock

Groundwater

10 n

5H2;aq þ SO¼ 4 , H2 Saq þ 4H2 O

8 6 4 2 0 –9

–6

–3

ð5Þ

δ34S (‰ VCDT)

Figure 3 | Stable isotope geochemistry of the biogenic rocks at Las Cruces. (a) Carbon–oxygen isotope composition of carbonates (sid, siderite; cc, calcite; ank, ankerite), overlying Tertiary sandstone and marl, and calculated d18O-d13C isotope composition of calcite in equilibrium with the local groundwaters (25 °C). Isotopic composition of the calcite in equilibrium with the water has been calculated from the tabulated d18O-d13C values together with the fractionation factors of Kim and O’Neil45 for oxygen and those of Romanek et al.46 for DIC. For comparison, data are shown for the Variscan (primary-ore related) hydrothermal siderite–ankerite assemblage occurring in VMS deposits of the Iberian Pyrite Belt. (b) Sulphur isotope composition compared with the d34Ssulphate values of the local groundwater and d34S values of the primary sulphides at Las Cruces19.

Microbes facilitating the classical reaction with methane and sulphate, producing CO2 and H2S (reaction 1), should be associated with another consortium of anaerobic bacteria that, in the presence of excess goethite, favour reaction (2) in which the goethite acts as an electron acceptor; this reaction is favoured due to its low free energy of formation28. Reactions (1) and (2) involve the oxidation of an electron donor, here represented by methane, but could well be organic molecules having low molecular weight. The low d13C values exclude hydrogen as the main electron donor and point instead to the involvement of methane or light hydrocarbons. The reduction of Fe3 þ to Fe2 þ associated with the replacement of goethite by siderite (reactions 2 and 3) is also probably due to bacterial activity. In surﬁcial low-temperature environments, the reduction of Fe3 þ to Fe2 þ occurs predominantly by biogenic processes29,30. Abiogenic reduction of iron is unlikely to take place at Las Cruces because in surﬁcial settings it has only been observed to occur in relationship with photoreduction31 and in some immature, organic matter-rich environments such as soils. In the latter setting, abiogenic reduction is associated with the oxidation of quinone-rich humic acids and phenolic compounds32, derived mainly from plant tissues. Even in the

ð6Þ

However, again the carbon isotope signatures indicate a significant input of carbon derived from the bacterial metabolization of methane/organic matter, therefore suggesting that reactions (1) and (2) are indeed relevant. The precipitation of pyrite seems to be kinetically inhibited in the Black Rock. Instead, amorphous or poorly ordered greigite or smythite have precipitated. Although these minerals have been formed abiotically in the laboratory33,34, Kucha and Barnes35 and Raiswell and Plant36 proposed that the presence of such intermediate-valence monosulphides is directly related to biogenic precipitation. At Las Cruces, crystallization of these iron sulphides inhibited the formation of siderite and rendered calcite the stable carbonate in the presence of high aCa2 þ /aH þ 2. There is no available Ca2 þ in the replaced gossan and, thus, we infer it is derived from the calcium-bearing groundwaters. Formation of the complex Ag-As-Fe-Cu sulphosalts intergrown with the galena is interpreted as being related to destabilization of the Ag-bearing jarosite and to input of sulphur of biogenic derivation. We also interpret the formation of free gold as being linked to bacterial activity: during sulphidization of the gossan, sub-microscopic ‘invisible’ gold11 was transformed into coarse grains; by analogy, multiple lines of evidences have been presented for the biogenic formation of gold nuggets37,38. Dissimilatory sulphate-reducing microbes have played a key role in the formation of many low-temperature stratiform ore deposits, with the reduced sulphur being derived from extensive biogenic reduction of dissolved sulphate22,39,40. Other sub-surface microbial activity is related to the secondary alteration of sulphides into oxides and carbonates41. During this supergene alteration, lithoautotrophic bacteria accelerate the dissolution of exposed ore deposits and promote the precipitation of secondary sulphides within the cementation zone, accelerating the process by up to five orders of magnitude above abiotic processes42,43. Biogenically mediated supergene precipitation of other sulphides has been described elsewhere, such as in the present-day formation of sphalerite biofilms in the Mike gold mine, in Nevada, USA23. We propose that the unusual mineral assemblage found in the secondary zone of the Las Cruces ore deposit formed in a large natural bioreactor, where coupled microbial sulphate reduction and methane oxidation took place below a thick impermeable marl unit. The preferred site for such biogenic activity is where sulphate-bearing groundwaters interacted with a previously formed porous gossan, which was also the zone of gas

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a SE

Buried gossan

Sandstone

Black Rock Cementation zone

Massive sulphides

Marl

NW Marl

Red Rock

Stockwork

Reworked gossan & sandstone Red & Black rock replacing former gossan

Biogenic zone

Sandstone

Red Rock Barren pyrite

Black Rock

Sulphate-bearing calcium-bicarbonate waters

Cementation zone

Massive sulphides

Upflowing methane or light hydrocarbons

Figure 4 | Proposed model for the genesis of the biogenic carbonate–sulphide rocks of the Las Cruces deposit. (a) General idealized crosssection of the deposit. The SE–NW distance is B300 m. (b) Cartoon based on the section depicted in a showing the proposed genetic model for the deposit. Carbonate–sulphide rocks (Red and Black rocks) formed in the supergene alteration zone (gossan) capping an older massive sulphide deposit. The porous gossan accumulated methane or light hydrocarbons that acted as electron donors for microbial reduction of aqueous sulphate transported by groundwater. This process destabilized goethite/haematite and oxidized minerals (beudantite, cerussite), leading to the formation of siderite, calcite, galena and silver-bearing sulphides, together with coarse-grained gold that has almost completely replaced the former gossan. Scale bar, B50 m.

Table 2 | Calculation of the amounts of sulphides produced by biogenic activity at Las Cruces. Grade Gossan Black Rock Calcite Galena Silver sulphides (as Ag2S) Gold Iron sulphides (as FeS) Red Rock Siderite Galena Haematite Total

1.7 Mt 40% 8.7%Pb 109 g/t Ag 3.85 g/t Au 60% 3.9%Pb

wt % 0.7 60.0 10.0 0.0 0.0 30.0 1.0 80.0 4.5 15.5

Tons

Moles metal

Moles CO2

Tons CO2

163,428 58,888 185 7 129,551

4.078E þ 09 2.842E þ 08 1.717E þ 06 3.322E þ 04 2.320E þ 09

4.078E þ 09

1.794E þ 05

326,857 39,750 110,596

6.779E þ 09 4.590E þ 08 1.581E þ 09

6.779E þ 09

1.086E þ 10

Moles H2S

Tons H2S

2.842E þ 08 8.587E þ 05

9.680E þ 03 2.925E þ 01

2.320E þ 09

7.902E þ 04

4.590E þ 08

1.563E þ 04

3.064E þ 09

1.044E þ 05

2.982E þ 05

4.776E þ 05

Galena has 86.6% Pb; thus, the total amount is 113,900 t. Ag2S has 87.07% Ag;thus, the total amount of silver sulphides as acanthite/argentite is 638 t. The total moles of calcite and siderite correspond to 1.19 Mt of total carbonates. The Black and Red rocks host B1.7 Mt of ore averaging 5.81% Pb, 3.85 g/t Au and 109 g/t Ag (www.ﬁrst-quantum.com); visual estimates in the pit and petrographic study suggest that 60% by weight of the capping zone is made of Red Rock (80% siderite, 15.5% haematite, 4.5% galena) and the remainder is Black Rock (60% calcite, 10% galena, 30% iron sulphides).

accumulation beneath the sealing marl (Fig. 4). Under these conditions, anaerobic sulphate-reducing microbes were able to reduce the aqueous sulphate that then reacted in situ with the available metals; in fact, the metal assemblage found within this biotic zone is the same as that found in the gossans of the Iberian Pyrite Belt11. This process would be synchronous with the release of CO2 as a byproduct of the heterotrophic metabolism of anaerobic microbes; mixing of this carbon with DIC transported by the groundwater produced supersaturation in, and precipitation of, carbonates. The inferred high CH4/SO4 ratios should promote quick and nearly complete sulphate reduction and the anomalously high sulphur isotope signatures. 6

An estimation of the tonnages and metal grades of the Red and Black rocks, based on mineral compositions and abundances, suggests that the sub-surface microbial system at Las Cruces has modified the mineralogy of at least 2–6 km3 of rock, fixing at least 1.09 1010 moles of CO2 and 3.1 109 moles of H2S, and resulting in the formation of ca. 1.19 Mt of siderite and calcite, 114,000 t of galena, 638 t of silver sulphides and 6.5 t of free gold (Table 2). Such quantities reflect the existence of a vast underground microbial system, with the limiting factor in this system probably being the availability of electron acceptors. The exact timing of the ore-forming process is difficult to constrain, because the assemblage lacks minerals suitable for

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accurate absolute dating. Geological relations suggest that the main gossan formation took place at 8–7 Ma (ref. 11) and that this gossan was covered by sediments at ca. 7.2–5.3 Ma; crosscutting relationships indicate that the deposition of galena postdated the formation of the underlying cementation zone and the overlying marl. Hence, the ore-forming event took place between burial at 7.2–5.3 Ma and the present. However, the microbially mediated mineralizing system has not changed significantly since burial and, thus, it may still be an active process. Our model for the formation of secondary Pb-Ag-Fe sulphides and gold is different from those invoked for all known biogenically mediated ore-bearing systems. Equivalent secondary mineral assemblages are virtually unknown in the Earth’s crust; the only broadly comparable assemblages that have been reported are in the small Zapadno–Ozernoe mine, in the Russian Urals44. The formation of the large bioreactor in the subsurface at the Las Cruces deposit probably involves a wide community of competing chemolithotrophic microbes resulting in mineralforming processes that are influenced and controlled by fluctuations in water input, temperature and availability of nutrients. Variations in any of these limiting factors and the nature of the aerobic/anaerobic interface probably control the size and resulting mineralogy of this vast underground ecosystem. Methods Sampling and preparation. Selected samples of ore, overlying sediments and waters were collected in situ from outcrops in the open pit; 1–3 g of individual minerals (carbonates and galena) were selected under the binocular microscope or separated magnetically (iron monosulphides); in the cases of finely intergrown assemblages, the gases were separated chemically following the procedures described below. Isotope measurements. Carbon dioxide was obtained by reaction of the carbonates47,48 and DIC49 with 100% phosphoric acid. Water was analysed using the CO2–H2O equilibration method of Epstein and Mayeda50. For analysis of dD, an aliquot of water (0.5 ml) was injected into a ceramic column containing a glassy carbon tube at 1,400 °C to produce H2 and CO gases51. Sulphur isotopes of dissolved sulphates were measured after precipitation of BaSO4 from solution. SO2 was obtained from the barium sulphate as well as from the sulphides by combustion with V2O5 and O2 at 1,030 °C (ref. 52). Isotope measurements were carried out at the Stable Isotope Laboratory of the Instituto Andaluz de Ciencias de la Tierra (CSIC-UGR, Granada) with a Delta Plus XL (ThermoQuest, Bremen, Germany) mass spectrometer (elemental analysis–isotope ratio mass spectrometry) for sulphur, oxygen in sulphate, and water and deuterium isotope analyses, and a Delta Plus XP spectrometer for C and O isotopes in carbonates. Precision was calculated, after correction of the mass spectrometer daily drift, from data obtained on standards systematically interspersed in analytical batches, at better than ±0.1% for d13C for carbonates and DIC, and for d18O carbonates and water oxygen. Precision was better than ±0.2% for sulphate sulphur (d34S) and for sulphate oxygen (d18O), and better than ±2% for water hydrogen (ddD). The standard for reporting carbon measurements is VPDB (Vienna-PDB), for oxygen and hydrogen is VSMOW (Vienna Standard Mean Oceanic Water) and for sulphur is VCDT (Vienna Canyon Diablo Troilite). Scanning electron microscope/energy-dispersive spectrometer analyses. Small, ca 2.5-mm grains of the Black Rock were analysed under a Jeol 5600-LV scanning electron microscope equipped with an Oxford Industries INCA X-sight energy-dispersive spectrometer at the Centro de Astrobiologıá. Backscattered and secondary electron images and energy-dispersive spectra were obtained on samples mounted on Al stubs and without coating (V ¼ 20 kV; I ¼ 85 mA, electron beam diameter ¼ 1 mm). Quantitative analysis performed on several (21) bacteriomorphic structures show that they are made of galena (mean 84.89±1.49% wt Pb, 15.11±1.02% wt S) with only negligible traces of other elements present.

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Acknowledgements This study was funded by project CGL2011-23207 of the SEIDI (Spain), the IPBSL project of the European Science Foundation and the ProMine project of the EU 7th Framework Program. We acknowledge the staff of Cobre Las Cruces, especially J.C. Baquero, I. Carrasco, C. Gomez, A. Francos and G. Obejero for granting access to the mine and for providing valuable data on geology and hydrogeology. We also thank R. Amils, C. Ayora, C. Conde, N.G. Miguelez, I. Sa´nchez-Andrea, J.C. Videira and N. White for fruitful discussions. Michael Russell offered thoughtful and constructive reviews of the original manuscript.

Author contributions F.T. and F.V. conceived of the study and together with J.M.E. did the ﬁeld work, and collected and prepared the samples. F.V. and C.M.-S. performed the electron microprobe and scanning electron microscope analyses, respectively; A.D. did the isotope analyses. F.T., J.F.S. and F.V. took the lead in writing the paper. All authors contributed to the discussion.

Additional information Competing ﬁnancial interests: The authors declare no competing ﬁnancial interests. Reprints and permission information is available online at http://npg.nature.com/ reprintsandpermissions/ How to cite this article: Tornos, F. et al. Formation of recent Pb-Ag-Au mineralization by potential sub-surface microbial activity. Nat. Commun. 5:4600 doi: 10.1038/ncomms5600 (2014).

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Terpenoids in extracts of Lower Cretaceous ambers from the Basque-Cantabrian Basin (El Soplao, Cantabria, Spain): Paleochemotaxonomic aspects César Menor-Salván a,*, Maria Najarro b, Francisco Velasco c, Idoia Rosales b, Fernando Tornos b, Bernd R.T. Simoneit d,e a

Centro de Astrobiología (CSIC-INTA), 28850 Torrejón de Ardoz, Spain Instituto Geológico y Minero de España (IGME), Rios Rosas 23, 28003 Madrid, Spain Universidad del País Vasco, Departamento Mineralogía y Petrología, Apdo. 644, 48080 Bilbao, Spain d COGER, King Saud University, Riyadh 11451, Saudi Arabia e Department of Chemistry, Oregon State University, Corvallis, OR 97331, USA b c

a r t i c l e

i n f o

Article history: Received 22 June 2010 Accepted 30 June 2010 Available online 3 July 2010

a b s t r a c t The composition of terpenoids from well preserved Cretaceous fossil resins and plant tissues from the amber bearing deposits of El Soplao and Reocín in Cantabria (northern Spain) have been analyzed using gas chromatography–mass spectrometry and the results are discussed using the terpenoid composition of extant conifers as a reference. Amber is present at many horizons within two units of coastal to shallow marine siliciclastics of Albian and Cenomanian age. The fossil resins are associated with black amber (jet) and abundant, well preserved plant cuticle compressions, especially those of the extinct conifer genus Frenelopsis (Cheirolepidiaceae). We report the molecular characterization of two types of amber with different botanical origins. One of them is characterized by the significant presence of phenolic terpenoids (ferruginol, totarol and hinokiol) and pimaric/isopimaric acids, as well as their diagenetic products. The presence of phenolic diterpenoids together with the lack of abietic and dehydroabietic acids excludes both Pinaceae and Araucariaceae as sources for this type of amber. The biological diterpenoid composition is similar to that observed for extant Cupressaceae. The second type of amber is characterized by the absence of phenolic terpenoids and other specific biomarkers. Some terpenoids with uncertain structure were detected, as well as the azulene derivative guaiazulene. Our results suggest that the amber from Cantabria could be fossilized resin from Frenelopsis and other undetermined botanical sources. The biological terpenoid assemblage confirms a chemosystematic relationship between Frenelopsis and modern Cupressaceae. Ó 2010 Elsevier Ltd. All rights reserved.

1. Introduction Amber is fossilized resin produced from the exudates of conifers and certain angiosperms and is considered to be one of the few fossil deposits of exceptional preservation (Konservat Lagerstätten), because it permits the conservation of fossil organisms with all their delicate anatomical details. Fossil resins not only preserve the anatomy of fossil life forms that were trapped as biological inclusions, but also constitute a valuable source of information about their own botanical origin, ancient terrigenous ecosystems and climatic change by means of their chemical composition (Anderson and Crelling, 1995).

* Corresponding author. Tel.: +34 91 520 6402 6458; fax: +34 91 520 1621. E-mail addresses: cmenor@amyp.es, menorsc@inta.es (C. Menor-Salván), m.na jarro@igme.es (M. Najarro), francisco.velasco@ehu.es (F. Velasco), i.rosales@igme.es (I. Rosales), f.tornos@igme.es (F. Tornos).

Analysis of the chemical composition of fossil resins is not straightforward, because the original biochemical fingerprints of the resins are usually modified during diagenesis, with the bioterpenoids (unmodified biosynthetic natural products) being transformed into geoterpenoids (diagenetic products of degraded bioterpenoids that are found in amber and fossil plant tissues; Otto et al., 2007). Despite these diagenetic alterations, geoterpenoids retain the basic skeletal structures of their biological precursors and can be used as molecular markers (biomarkers; Peters et al., 2005; Marynowski et al., 2007). Conifers synthesize mainly diterpenoids, which are, along with sesquiterpenoids, the compounds that provide the best results as diagnostic biomarkers of conifers and their resins (Otto and Wilde, 2001). Among the diterpenoids preserved in amber, labdane derivatives and non-phenolic abietane diagenetic derivatives have the most limited chemotaxonomic value, as they occur in all conifer families. On the other hand, phenolic terpenoids, such as ferruginol and totarol, are produced only by

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the members of the families Cupressaceae and Podocarpaceae (Cox et al., 2007). Therefore, the chemotaxonomic value of these compounds is very high and their presence in amber provides very useful palaeobotanical information. Although the preservation potential of polar biomarkers is considered to be low (Otto et al., 2007), the oldest polar diterpenoids have been identified in extracts of Middle Jurassic fossil conifer wood from Poland (Marynowski et al., 2007), and diterpenoid derivatives could also be liberated from a Carboniferous amber by pyrolysis (Bray and Anderson, 2009). Recently, a new Cretaceous amber deposit with exquisite, well preserved fossil organisms, mostly insects, has been discovered in northern Spain (Rábago village in El Soplao territory, Cantabria; Menor-Salván et al., 2009; Najarro et al., 2009). Based on preliminary infrared spectroscopy of the El Soplao amber (Najarro et al., 2009) and previous gas chromatography–mass spectrometry (GC–MS) studies on amber from a neighboring site in Álava (Alonso et al., 2000; Chaler and Grimalt, 2005), it has been suggested that exudate from Agathis (a conifer of the family Araucariaceae) was the most likely source of this amber, as has also been proposed for other Cretaceous ambers (e.g. Lambert et al., 1996; Alonso et al., 2000; Poinar et al., 2004; Chaler and Grimalt, 2005; Delclòs et al., 2007). This speculation was largely based on the presence of some geoterpenoids that may have been derived from agathic and pimaric acids. However, although those compounds and their diagenetic derivatives are characteristic of Araucariaceae, they are not diagnostic, because they can also be found in other extant conifer families (Otto et al., 2007). Moreover, Alonso et al. (2000) have reported the presence of the phenolic abietane ferruginol in the Álava amber samples, indicating that more extensive study of the chemotaxonomic information contained in the amber is necessary to establish its definite botanical origin. In addition, meso- and macrofossil plant remains of Araucariaceae are absent in these amber bearing deposits, although there are plenty of cuticles and remains of other vascular plants, especially the genera Frenelopsis sp. and Mirovia sp., of the extinct conifer families Cheirolepidiaceae and Miroviaceae, respectively (Gomez et al., 2002a; Najarro et al., 2009). This is also the case in many other amber deposits from the Cretaceous of Spain and France (e.g. Delclòs et al., 2007; Néraudeau et al., 2008). Thus, the recurrent association of amber with cuticles of Cheirolepidiaceae and Miroviaceae, along with the lack of Araucariaceae remains (except for a small amount of pollen grains in the sediments) (Barrón et al., 2001), challenges the proposed origin of the amber. Since chemical evidence has not yet given a definitive answer, more convincing proof is necessary to accept Araucariaceae as the source of the resin. In this study, amber pieces and associated fossil leaves from the Cretaceous amber bearing deposit of El Soplao (Cantabria; Fig. 1) were systematically analyzed using complementary techniques such as infrared spectroscopy (FTIR) and GC–MS. The overall aim was to identify the terpenoids preserved and their diagenetic transformation products in the fossil resin and to determine their possible botanical sources. Due to exceptional preservation, the amber bearing deposit at El Soplao offers a unique opportunity to compare the molecular composition of the amber with that of plant remains from the family Cheirolepidiaceae and Miroviaceae, which appear in the same deposit. A morphological similarity between extinct Cheirolepidiaceae and extant Cupressaceae has been described, but their relationship remains speculative due mainly to the lack of molecular evidence (Broutin and Pons, 1975; Alvin and Hluštík, 1979; Seoane, 1998; Miller, 1999; Farjon, 2008). As Cheirolepidiaceae is an extinct family, the connection between the two families could aid in the chemotaxonomical study of amber and in the confirmation of the botanical origin. We present data of two separate types of amber found in the El Soplao deposit and discuss their botanical origin using comparative chemotaxon-

omy based on modern resin compositions and related terpenoids found in amber associated fossils.

2. Samples and methods 2.1. Geological background The analyzed samples belong to the Cretaceous succession at the northwestern margin of the Basque-Cantabria Basin in northern Spain. During the Cretaceous, the evolution of this basin was controlled by extensional, and perhaps strike-slip, deformation associated with the opening of the North Atlantic Ocean and the Bay of Biscay (e.g. Le Pichon and Sibuet, 1971; Rat, 1988; García-Mondéjar et al., 1996; Soto et al., 2007). Rifting during the Late Jurassic–Early Cretaceous led to the formation of several narrow sub-basins controlled by E–W, NW–SE and SW–NW trending faults; these basins host both continental and marine sediments of variable thickness (García-Mondéjar et al., 1996; Soto et al., 2007). The study area lies in the Cantabria region immediately to the north of the Cabuérniga Ridge (Fig. S1; supplementary material), an E–W fault zone that represents a Late-Variscan structure reactivated first as a paleo-high bounded by normal faults during the Early Cretaceous, and later as reversal faults during the widespread Cenozoic (Pyrenean) compression. The Lower–Middle Cretaceous (Barremian–Early Cenomanian) deposits in the study area are weakly deformed and affected only by gentle folding. They are composed of a relatively thin ( 200–800 m) syn-rift sequence that lies unconformably on Carboniferous to Lower Jurassic basement (Fig. S1). A simplified synthesis of the stratigraphy in the El Soplao and Reocín areas is shown in Fig. 1, with formation names according to Hines (1985) and revised by Najarro et al. (2009). The amber bearing deposit at El Soplao is included within the Las Peñosas Formation (Fig. 1), a Lower Albian unit ( 112–110 Ma) of continental to transitional marine siliciclastics. Detailed descriptions of field sections, depositional environments and fossil content of this unit are given in Najarro et al. (2009). Within the outcrop, the El Soplao amber deposit is characterized by about 1.5–2 m of dark, carbonaceous lutites, siltstones and sandstones with interbedded, centimeter to decimeter layers with remarkable accumulations of plant remains and amber pieces of different sizes and forms (Fig. 2A and B). Most amber pieces show a blue-purple color under normal sunlight and bright milky blue fluorescence under ultraviolet light. Plant cuticles are very abundant in the levels associated with amber (Fig. 2C). They are mainly assigned to the conifer genera Frenelopsis and Mirovia, along with other more occasional leaves of the ginkgoalean genera Nehvizdya and Pseudotorellia (Najarro et al., 2009). In most of the amber beds, leaves of the genus Frenelopsis of the extinct conifer family Cheirolepidiaceae are the dominant macro-botanical remains. Frenelopsis were xeromorphic plants adapted to coastal habitats and probably grew mainly in brackish coastal marshes and mangroves, but were adapted to a wider range of habitats (Gomez et al., 2002a, 2003). The amber deposit at Reocín (Fig. S1; supplementary material) is slightly younger than the El Soplao amber deposit. It is included within the Bielva Formation (Fig. 1), a Latest Albian–Early Cenomanian ( 102–99 Ma) unit composed of about 250 m of tidal dominated, estuarine siliciclastic deposits in the study area (Hines, 1985). Within this unit, the amber accumulations are associated with carbonaceous claystones and tidal channel sandstones developed in estuarine mouth subtidal areas (López-Horgue et al., 2001). Despite the differences in age, the Reocín amber shows the same composition as the El Soplao amber. Thermal maturity indicators (vitrinite reflectance) of macerals in the El Soplao deposit reveal minor changes in the organic matter of the resins during their diagenetic history and maximum thermal conditions during burial of

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SW LITHOLOGY

145 ma

200 ma

250 ma

300 ma

360 ma

LATE EARLY

CEN ALBIAN

BARCENACIONES Fm. LAS PEÑOSAS Fm.

LATE

APTIAN

LITHOLOGY

ALTAMIRA Fm. BIELBA Fm.

REOCÍN Fm.

EARLY

125 ma

STRATIGRAPHY OF REOCÍN SECTOR

BARCENACIONES Fm. LAS PEÑOSAS Fm.

SAN ESTEBAN Fm. PATROCINIO Fm. UMBRERA Fm. RÁBAGO Fm. PAS GROUP (WEALD)

BAR HAU VAL BER

CARBO PER TRIA JURA NIFE MIAN SSIC SSIC ROUS

112 ma

CRETACEOUS

99 ma

STRATIGRAPHY OF EL SOPLAO SECTOR

EAR LAT

93 ma

REOCÍN Fm. RODEZAS Fm.

SAN ESTEBAN Fm. PATROCINIO Fm. UMBRERA Fm.

PAS GROUP. (WEALD)

HIATUS

LIAS KEUPER BUNTSANDSTEIN

BUNTSANDSTEIN

HIATUS

PALEOZOIC BASEMENT PALEOZOIC BASEMENT Marls

Mudstones and gypsum

Conglomerates

Limestones

Mudstones and sandstones

Sandstones

Dolostones

Mudstones

Silts

Amber-bearing deposit

Lignite

Fig. 1. Chrono- and lithostratigraphy of the El Soplao and Reocín areas (modiﬁed from Hines, 1985). Chronostratigraphy after Gradstein et al. (2004).

60–70 °C (Supplementary material). Consequently and due to its higher transparency and lack of inclusions and interferences, the El Soplao amber was used preferentially for the chemosystematic study. 2.2. Sampling Amber pieces, jet (black amber), fossil wood and sediments rich in plant cuticles were collected from the El Soplao deposit during a recent excavation in October 2008. Two types of amber pieces were found at the deposit in the same sedimentological and taphonomical context: A type, characterized by a strong blue-purple color under natural light, purple-reddish under artiﬁcial light and less abundant B type, yellow-honey under artiﬁcial light and honey with a bluish tinge under natural light. We collected the two types of amber present and the black amber associated with amber of type A and fossil plant tissue. Plant cuticles were obtained from claystones by rinsing the plant rich sediment in an ultrasonic bath of distilled water to remove all the clay and silt sediment. The organic residue (Fig. 2C) was air dried. Plant fragments and leaves from different families were distinguished and separated under a stereomicroscope. 2.3. Analytical methods 2.3.1. Infrared spectroscopy (FTIR) IR spectra of pulverized solid amber were obtained using a Nexus Nicolet FTIR spectrometer in the 4000–400 cm 1 range.

2.3.2. Extraction and fractionation For the analytical characterization, two representative single pieces (A and B; Table 1) of amber of about 50 g, with the highest transparency available and free of major inclusions, crusts and debris, were collected from the El Soplao deposit. Each piece was crushed and extracted for 4 h with dichloromethane:methanol (2:1 v:v) using a Büchi model B-811 automatic extractor. The extractable material constitutes 16% of the total amber weight on average. One aliquot of extract was injected directly into the port of the gas chromatograph. The bulk extract was then processed in order to purify the phenolic terpenoid fraction and the acidic fraction and to identify unambiguously the minor components with higher chemosystematic value. The aim was to establish a complete descriptive composition of the amber sample. The extract was concentrated to a volume of 20 ml and fractionated by ﬂash chromatography on silica gel. The elution was performed using n-hexane, dichloromethane, dichloromethane:methanol (1:1 v:v), and methanol as eluents and 25 fractions of 1.5 ml were collected using an automatic fraction collector. Each fraction was concentrated by evaporation of the solvent under N2 and analyzed by GC–MS. The fractions with similar compositions were combined. The polar fraction (eluted with methanol) and the fractions containing ferruginol were recombined, further separated using a glass column (20 cm) ﬁlled with chromatographic grade silica gel, and eluted sequentially with n-hexane:dichloromethane (1:1 v:v), pure dichloromethane, dichloromethane:methanol (1:1 v:v) and methanol. Four fractions of 20 ml were collected, designated A to D. All fractions were dried and the alcohols and acids converted to trimethylsilyl derivatives

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Fig. 2. Photographs of amber and palaeobotanical components used in this study. (A) In situ amber piece of type A (6 cm) from the El Soplao deposit, showing blue-purple color under sunlight. (B) Amber piece of type B (3 cm) on fossil wood under sunlight. (C) In situ sediments showing their palaeobotanical components: fossil leaves of Mirovia and Frenelopsis and one leaf of Ginkgoaceae. (D) Selected leaves of Frenelopsis sp. were used for biomarker analysis. (For interpretation of the references to color in this ﬁgure legend, the reader is referred to the web version of this article.)

by reaction with N,O-bis-(trimethylsilyl)trifluoroacetamide (BSTFA) containing 1% trimethylchlorosilane (TMCS) at 65 °C for a period of 3 h. Finally, the derivatized fractions were diluted with n-hexane and injected into the port of the gas chromatograph. To study the molecular content of fossil Frenelopsis and Mirovia leaves (Fig. 2), 5 g of leaves were extracted for 4 h with dichloromethane:methanol (2:1 v:v) using a Büchi model B-811 automated extractor. The bulk extract was filtered and analyzed directly by GC–MS. The extract was then fractionated using silica gel chromatography in two fractions by elution with n-hexane:dichloromethane (3:1 v:v) and dichloromethane:methanol (4:1 v:v). The polar fraction was dried and derivatized using the method described above. The samples of jet (black amber) were extracted using the same protocol. Due to the lesser availability of leaves and jet and the lower percentage of extractable organic matter, we used the simplified fractionation described above in order to compare their biomarker composition with that of amber. 2.3.3. GC–MS The analyses were performed on an Agilent 6850 GC coupled to an Agilent 5975C quadrupole mass spectrometer. Separation was achieved on a HP-5MS column coated with (5%-phenyl)-methylpolysiloxane (30 m 0.25 mm, 0.25 lm film thickness). The operating conditions were as follows: 8 psi carrier pressure, initial temperature held at 40 °C for 1.5 min, increased from 40 °C to 150 °C at a rate of 15 °C/min, held for 2 min, increased from 150 °C to 255 °C at a rate of 5 °C/min, held constant for 20 min and finally increased to 300 °C at a rate of 5 °C/min. The sample was injected in the splitless mode with the injector temperature at 290 °C. The mass spectrometer was operated in the electron impact mode at 70 eV ionization energy and scanned from 40 to 700 Da. The temperature of the ion source was 230 °C and the quadrupole temperature was 150 °C. Data were acquired and processed using Chemstation software. Individual compounds were

identiﬁed by comparing their mass spectra with those of authentic standards and with published data (see Section 3.2). 3. Results and discussion 3.1. Infrared spectroscopy The application of IR to the study of amber is well documented and constitutes a basic technique for the characterization of fossil resins (Langenheim, 1969; Grimalt et al., 1988; Alonso et al., 2000). Because of the inclination of all ambers and resins (even non-fossil resins) to show similar bulk infrared spectra (due to their common chemical functional groups), IR spectroscopy has strong limitations for the determination of their botanical origin (Yamamoto et al., 2006). The IR spectrum of the Cantabrian amber is consistent with those observed for other amber samples (Fig. S2; supplementary material) and could indicate that it is composed of a mixture of terpenoids and labdatriene copolymers. The band pattern is similar to the IR spectrum expected for the labdatrienes communic acid and biformene and their polymers, consistent with the stated macromolecular structure of amber (Villanueva-García et al., 2005) and with the terpenoid composition found (see Section 3.2). The weak band at 882 cm 1 (Fig. S2, supplementary material) is characteristic of the exocyclic methylene moiety supporting the labdatriene input. The two types of amber samples found in the deposit show similar IR spectra. 3.2. Terpenoid composition of Cantabrian amber The total extracts of the amber contain methylated naphthalenes (di-, tri- and tetramethylnaphthalenes) and di- and trimethyltetralins, sesquiterpenoids and bi- and tricyclic diterpenoids (Table 1). GC analysis of the bulk extract shows three different zones in the gas chromatogram (Fig. 3). The dominant compounds in the early

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Compound

Composition

Relative abundanceb A

224 228 228 234 242 252 256 256 256 256 258 260 262 268 270 272 286 300 302

38.2 100 23.4 1.0 4.6 9.3 2.5 – 45.8 3.6 9.7 57.0 5.8 11.8 40.3 d 24.3 d d

– 100 20.9 – 42.5 1.5 6.9 6.6 14.0 4.1 – 23.0 d – – d – – –

C20H30O2 C20H30O2 C20H32O2 C20H32

302 302 304 272

d d d –

d d d 4.5

14,15-Bisnorlabda-8,12-dien-18-oic acidc E-19-Noragathic acid Z-19-Noragathic acid 13-Dihydro-19-noragathic acidc 13-Dihydroagatholic acid

C18H28O2 C19H30O2 C19H30O2 C19H32O2 C20H34O3

276 290 290 290 322

d 10.6 3.7 d 11.8

d 3.1 1.5 – –

Ionene Methylionene Tetrahydroeudalene Guaiazulene Cadalene Drimane Homodrimane 2,5,8-Trimethyl-1-butyltetralin 2,5,8-Trimethyl-1-isopentyltetralin Diaromatic totarane Totarolc

C13H18 C14H20 C14H20 C15H18 C15H18 C15H28 C16H30 C17H26 C18H28 C19H24 C20H30O

174 188 188 198 198 208 222 230 244 252 286

37.2 13.0 16.0 7.6 8.2 3.6 3.3 79.1 d 1.0 d

48.4 34.2 8.3 1.5 7.5 14.8 3.9 57.3 10.8 – –

Abietanes and Podocarpanes 1 (XVI) 16,17-Bisnorsimonellite 2 (XXII) 16,17,18-Trisnorabieta-8,11,13-triene 3 16,17,19-Trisnorabieta-8,11,13-triene 4 (XVII) Retene 5 16,17-Bisnordehydroabietane 6 Simonellite 7 (XXVII) 14-Methyl-16,17-bisnordehydroabietane 8 1-Methyl-10,18-bisnorabieta-8,11,13-triene 9 (I) 18-Norabietatriene (Dehydroabietin) 10 (I) 19-Norabietatriene 11 (IV) 18-Norabieta-7,13-diene 12 (V) Norabiet-13-ene 13 (III) Fichtelite 14 (X) 12-Hydroxysimonellite 15 (II) Dehydroabietane 16 (XXV) 16,17-Bisnordehydroabietic acidc 17 (VIII) Ferruginol 18 (XIII) Callitrisic acidc 19 (XII) Hinokiolc

C17H20 C17H24 C17H24 C18H18 C18H26 C19H24 C19H28 C19H28 C19H28 C19H28 C19H30 C19H32 C19H34 C19H24O C20H30 C18H24O2 C20H30O C20H28O2 C20H30O2

Pimaranes and Isopimaranes 20 (XXI) Pimaric acidc 21 (XXII) Isopimaric acidc 22 (XXVI) Pimar-8-en-18-oic acidc 23 Pimaradiene Labdanes 24 25 (XXXIV) 26 (XXXIII) 27 (XXIX) 28 (XXVIII) Other compounds 29 30 31 32 (XXXVII) 33 (XXXVIII) 34 35 36 (XXXI) 37 (XXXII) 38 (XX) 39 (XI) a

Roman numerals in parentheses refer to structures shown in Appendix A. Abundance relative to the major peak (100%) in the bulk extracts (GC–MS TIC). Occurrence is tabulated on compounds detected only after fractionation and derivatization (d: detected; –: not detected). c Also analyzed as the TMS derivative. b

elution range are a-ionene, methylionene, trimethylnaphthalene isomers, tetrahydroeudalene, calamenene isomers, drimane and homodrimane (identified after Dzou et al., 1999 and Sonibare and Ekweozor, 2004). The a-ionene, methylionene and drimanes may be derived from labdanes in the resin through degradation processes (Yamamoto et al., 2006; Pereira et al., 2009). Overall, these components are highly degraded diagenetic products that have no chemotaxonomic value due to their unrecognizable parent structures. The second section of the gas chromatogram is dominated by non-oxygenated bi- and tricyclic diterpenoids and the third section contains polar bi- and tricyclic diterpenoids. We did not find aliphatic lipids, hopanoids, fungal terpenoids or plant triterpenoids in the amber samples, discarding an angiosperm contribution and major contamination. 3.2.1. Abietane diterpenoids The diterpenoids identified in the extracts belong to the abietane, pimarane/isopimarane and labdane structural classes (Fig. 3). These diterpenoids are typical of conifers (Otto and Wilde, 2001; Yamamoto et al., 2006), confirming such an origin for the Cantabrian ambers.

The abietane class terpenoids were identified by comparison of their mass spectra with those of standards or published in the literature (Czechowski et al., 1996; Otto and Simoneit, 2002; Otto et al., 2002; Hautevelle et al., 2006; Cox et al., 2007), and comprised 18- and 19-norabieta-8,11,13-triene (I; chemical structures cited are shown in Appendix A), dehydroabietane (II), fichtelite (III), 18-norabieta-7,13-diene (IV) and norabiet-13-ene (V). The latter compound was tentatively identified by match with a mass spectrum in the literature (Hautevelle et al., 2006), characterized by a molecular ion at m/z 260 and loss of an isopropyl group (m/z 217). 18-Norabieta-7,13-diene (IV) was identified only in sample A by a match with the published mass spectrum (Otto and Simoneit, 2002). This compound has been described as a decarboxylation product of abietic acid during diagenesis (Otto and Simoneit, 2002). In this case, the precursor molecule has not been found. The lack of a clear biological precursor for norabieta-7,13-diene (IV) suggests an alternative origin, possibly by double bond isomerization of unsaturated abietanes. This composition is consistent with the dominance of dehydroabietane and abietane geoterpenoids in the type A amber. The norabietatrienes (dehydroabietins) found in both amber

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A Dehydroabietin

16,17,18-Trisnorabieta- 18-Norabieta -7,13- Norabiet -13-ene 8,11,13-triene diene

OH O

16,17Bisnordehydroabietane

16,17Bisnorsimonellite

2,5,6-trimethyl1-butyl tetralin

19-Noragathic acid

Tetrahydroeudalene

Relative abundance

Dehydroabietane

Guaiazulene α-Ionene

13-Dihydroagatholic acid

Ferruginol

Methylionene

Homodrimane

Simonellite 12hydroxysimonellite

Drimane

Methylionene

2,5,6-Trimethyl-1isopentyl tetralin

16,17,18-Trisnorabieta8,11,13-triene

Norabiet -13-ene Pimaradiene α-Ionene

Dehydroabietin

2,5,6-Trimethyl1-butyl tetralin

Relative abundance

16,17Bisnordehydroabietane

Guaiazulene

Drimane

19-Norlabda 8(20),12-dien -15-oic acids

14-Methyl -16,17bisnordehydroabietane

OH O

Retention time (min) Fig. 3. GC–MS total ion current (TIC) traces of the underivatized total extracts of: (A) El Soplao blue amber, type A, (B) El Soplao yellow-blue amber, type B, with the main terpenoid compounds identified. Peaks not annotated are unidentified, tentatively identified or known compounds with little chemotaxonomic value. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

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samples could be derived from all abietane precursors by diagenetic alteration (Simoneit, 1986; Hautevelle et al., 2006). Dehydroabietane (II) is a natural product of many Pinaceae resins (Otto et al., 2007) as well as some Cupressaceae resins. In these samples, dehydroabietane is a significant component in amber type A, whereas it is not detectable in amber type B, suggesting a different paleobotanical origin for both types of amber samples. The absence of abietic (VI) or dehydroabietic acids (VII) eliminates a Pinaceae contribution to the amber, because abietic acid is a major component of such resins and dehydroabietic acid, its major diagenetic derivative, is present in ambers derived from Pinaceae (Yamamoto et al., 2006). Phenolic diterpenoids occur in polar fraction B of the type A amber (Fig. 4), with a dominance of ferruginol (VIII) and its oxidation products 6,7-dehydroferruginol (IX) and 12-hydroxysimonellite (X) and totarol (XI) (Otto and Simoneit, 2001; Otto et al., 2002). Ferruginol and 12-hydroxysimonellite are also identifiable (underivatized) in the bulk extract as part of the main components of the amber (Fig. 3). The presence of these phenolic diterpenoids is of significant chemosystematic value, as ferruginol is an abundant natural product in extant conifers of the families Cupressaceae and Podocarpaceae and can be used as a characteristic biomarker of these families (Otto and Simoneit, 2001; Marynowski et al., 2007). A minor amount of hinokiol (3-hydroxyferruginol, XII; Fig. 5A) was also found in the type A amber. Hinokiol has been described from Cupressaceae (Otto et al., 2002; Cox et al., 2007). There is no reported presence of phenolic diterpenoids in modern Araucariaceae (Cox et al., 2007), but Otto and Wilde (2001) cited the occurrence of ferruginol in Araucaria. To avoid this ambiguity and to test this finding under our experimental conditions, resins of Agathis sp. and Araucaria sp. were analyzed and ferruginol was not detected in any Araucariaceae resins. Hence, these results, coupled with the absence of kaurane or phyllocladane diterpenoids, the

Ferruginol

presence of totarol (see below) and the fossil record of the deposit, suggest that Araucariaceae did not contribute to the main type of amber found at the studied deposit (type A). On the other hand, we did not find phenolic abietanes in the type B amber sample. This fact, taken together with the presence of dehydroabietane in sample A and its absence in sample B, constitutes the main chemotaxonomic difference between the two types of samples. A significant relationship between amber type A and modern Cupressaceae is the presence of a low quantity of callitrisic acid (XIII) which is an epimer of dehydroabietic acid (VII, Fig. 5). The difference in retention time with dehydroabietic acid and the higher relative intensity of the ion at m/z 357 (M-CH3) versus the molecular ion in callitrisic acid are distinctive features between the two epimers used here for the identification of the acid (Van den Berg et al., 2000; Cox et al., 2007). Callitrisic acid has a higher chemotaxonomical value than dehydroabietic acid due to its scarcity. In modern conifer resins, the synthesis of callitrisic acid seems to be restricted to certain genera of the Cupressaceae family and it was also found in Cenomanian amber from the Raritan Formation (New Jersey, USA), suggesting a relationship with Cupressaceae (Anderson, 2006). Degradation of phenolic diterpenoids could lead to the abietane geoterpenoids found in the type A amber (Otto et al., 1997; Otto and Simoneit, 2001; Stefanova et al., 2002). Hautevelle et al. (2006) and Yamamoto et al. (2006) discussed the diagenetic pathways of abietane class bioterpenoids, suggesting that 18-norferruginol (XIV) could be the precursor of dehydroabietin, and ferruginol (VIII) could lead to 12-hydroxysimonellite (X), simonellite (XV), 16,17-bisnorsimonellite (XVI) and retene (XVII), all found in the type A amber. Under the anaerobic depositional conditions of the amber (Najarro et al., 2009), we could not disregard redox reactions that lead to the actual composition found (Pereira et al., 2009). If, as in some modern Cupressaceae genera (i.e. Cupressus; Fig. S3, supplementary material), the original proportion of ferruginol (VIII) was high, its

12Hydroxysimonellite

Relative abundance

OH 6,7-Dehydroferruginol

Totarol

Retention time Fig. 4. GC–MS TIC trace of fraction B resulting from the column separation of the polar fraction of amber A. This fraction contains the partially puriﬁed phenolic terpenoids. Compounds are analyzed as the TMS derivatives.

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Fig. 5. GC–MS TIC traces of the polar fractions C of: (A) amber A and (B) amber B. This fraction contains mainly pimaric and labdenoic acids. The figure shows the identified compounds in the fraction. Unlabelled peaks are considered as unidentified due to the lack of standards and published references or databases with details of their mass spectra. Compounds are analyzed as the TMS derivatives.

diagenesis could ultimately have generated dehydroabietane (II) and simonellite (XV), both significant in the type A amber. 3.2.2. Totarol Totarol (XI), a tricyclic diterpenoid phenol, is considered as a confirmatory chemotaxonomic marker for Cupressaceae and Podocarpaceae, even at low concentrations (Le Métayer et al., 2008; Stefanova and Simoneit, 2008). Totarol is detectable in the bulk extract using the characteristic mass fragments of m/z 271 and 286. The identification is unambiguous in amber sample A after purification of the phe-

nolic fraction of the bulk extract and analysis as trimethylsilyl derivatives (Fig. 4). Due to the similarities between the mass spectra of phenolic diterpenoids, the retention time and mass spectrum of totarol (XI) were determined using a standard (Sigma–Aldrich). The presence of totarol suggests a relationship between the palaeobotanic origin of the amber and extant Cupressaceae or Podocarpaceae. To test this possibility, the phenolic diterpenoids of the amber were compared with those from a modern Cupressaceae (Cupressus arizonica; Fig. S3, supplementary material). Both extracts contain ferruginol (VIII), totarol (XI) and hinokiol (XII) as the main phenolic

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diterpenoids, but sempervirol (XVIII) has not been observed. A difference between the assemblage of polar terpenoids from C. arizonica and the amber is the presence of sugiol (XIX) and the lack of callitrisic acid (XIII) in the former. We identiﬁed a diaromatic totarane (XX) as a possible diagenetic product of totarol (XI), which may be derived by a parallel diagenetic pathway as simonellite (XV) from ferruginol (VIII) (Otto et al., 1997). In accord with the phenolic abie-

tane composition, totarol is not detectable in the type B amber, conﬁrming that both amber types found in the Cantabria deposits differ in their biological origins. 3.2.3. Pimarane/isopimarane diterpenoids Polar fraction C of amber contains low amounts of pimaric (XXI) and isopimaric (XXII) acids (Fig. 6). Diagenesis of pimarane

6,7-Dehydroferruginol

Ferruginol

Dehydroabietane

Dehydroabietin

12-Hydroxysimonellite

Tetrahydroretene

Simonellite

16,17-Bisnorsimonellite

? Pimaradiene

14-Methyl-16,17bisnordehydroabietane

Isopimaradiene

16,17Bisnordehydroabietane

OH O

Pimaric acid Isopimaric acid

Pimar-8-en-18-oic acid

16,17Bisnordehydroabietic acid

16,17,18-trisnorabieta8,11,13-triene

HO O

E- and Z-19-norlabda-8(20),12-dien-15-oic acids

Agathic acid 13-Dihydroagatholic acid 2,5,8-Trimethyl-1-alkyltetralins Fig. 6. Proposed diagenetic pathways for the diterpenoid precursors from the Cantabrian ambers (based on Otto and Simoneit (2002), Stefanova et al. (2002), Hautevelle et al. (2006), and Pereira et al. (2009)). Dotted box: biological precursors; solid box: major terpenoid found in the samples.

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A 16,17Bisnorsimonellite

Tetrahydroretene

Dehydroabietane

Dehydroabietin 16,17,18-Trisnorabieta8,11,13-triene

Simonellite

14-methyl-16,17bisnordehydroabietane

Pristane

Relative abundance

2,5,8-Trimethyl1-butyl tetralin

Cadalene Retene

Ferruginol

Naphthalene and tetralin derivatives

12-Hydroxysimonellite

16,17,18-Trisnorabieta8,11,13-triene

16,17Bisnorsimonellite

Cadalene

Dehydroabietane

Simonellite 2-Methylretene

Relativeabundance

Guaiazulene OH

Retene

Naphthalene and tetralin derivatives

12Hydroxysimonellite

Ferruginol

Retention time (min) Fig. 7. GC–MS TIC traces of the total extracts from: (A) Frenelopsis leaves and (B) jet (black amber), showing the identiﬁed biomarkers.

diterpenoids could be one of the possible origins of 16,17,19-trisnorabieta-8,11,13-triene (XXIII), a major compound identiﬁed in the amber samples. Another possible origin for this compound is by diagenesis of dehydroabietane and other abietane related terpe-

noids (Fig. 6; Otto et al., 2002; Pereira et al., 2009). Due to the widespread distribution of the pimaric/isopimaric acids, the chemotaxonomical interpretation of their presence in the Cantabrian ambers must be taken with caution and comparisons with

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other markers and samples should be made. Continuing the comparison with extant conifers, pimaric and isopimaric acids constitute the main tricyclic resin acids present in the conifer families Cupressaceae, Araucariaceae and Podocarpaceae (Otto and Wilde, 2001). Taking into account a possible molecular relationship between the botanical source of Cantabrian amber A and modern Cupressaceae, if pimaric acid (XXI) and isopimaric acid (XXII) were the main resin acids in the precursor resin of these ambers, diagenetic degradation to 16,17,18-trisnorabieta-8,11,13-triene (XXIII) is consistent with the dominance of the pimarane resin acid compounds in the extract, as loss of the vinyl moiety at C-13, with concomitant aromatization and decarboxylation at C-4 generates this predominant isomer. Following this pathway, the presence of a related molecule to 16,17,18-trisnorabieta-8,11,13-triene with a C-13 ethyl group (i.e. 16,18-bisnorabieta-8,11,13-triene, XXIV) should be expected as well, but we failed to detect such a compound. Otto et al. (2002) reported a significant presence of that geoterpenoid (XXIV) in fossil resin of the Lower Cretaceous Tritaenia linkii (Miroviaceae), but due to the absence of precursor bioterpenoids, the assignment to a specific taxon was unclear. Recent work of Pereira et al. (2009) on Cretaceous amber from Brazil, inferred some intermediates of this diagenetic route (Fig. 6), namely 16,17-bisnordehydroabietic acid (XXV), pimar-8-en-18-oic acid (XXVI) and 14-methyl-16,17-bisnordehydroabietane (XXVII), that are also present in the type A amber (Figs. 3–5). Work is in progress in our laboratory in order to confirm this hypothetical pathway, as the formation of 14-methyl-16,17-bisnordehydroabietane by rearrangement of a pimaradiene or abietane precursor has not been demonstrated to date. 3.2.4. Labdane diterpenoids Labdanoic acids and other labdane derivatives are common components in all conifers and are therefore non-specific biomarkers (Otto and Wilde, 2001). 13-Dihydroagatholic acid (XXVIII) is the predominant labdenoic acid present in the polar fraction of amber A (Fig. 5). This acid could be a precursor molecule preserved that constitutes the chemotaxonomic difference between the two paleobotanical resin producers, as it is not detectable in the extract of amber B. 13-Dihydro-19-noragathic acid (XXIX), found in sample A, could be formed from the 13-dihydroagatholic acid (XXVIII) precursor by loss of the C-19 hydroxymethyl group at C-4 or from the agathic acid (XXX) precursor by C-19 decarboxylation at C-4, respectively. The diagenetic transformation of the major labdanoic acids may be the source of the 2,5,8-trimethyl-1-alkyltetralins, ionenes and drimanes found in the samples (Fig. 3). The MS fragmentation pattern of the major compound of this family, i.e. 2,5,8-trimethyl-1-butyltetralin (XXXI) with a molecular ion at m/z 230, shows a butyl loss (57 da) from the saturated ring to form an m/z 173 fragment (see Fig. S4, supplementary material). Another homologue of this compound group is 2,5,8-trimethyl-1isopentyltetralin (XXXII), which is significant in amber type B but only occurs in trace amounts in amber A. The degradation pathway leading to XXXII could be decarboxylation at C-4 of a labdanoic acid precursor (i.e. agathic acid, XXX), followed by aromatization of ring A with methyl migration from C-10 to C-1 and decarboxylation of C-15 (Fig. 6). These compounds were also reported from Brazilian ambers (Pereira et al., 2009). The difference found between the ambers could be indicative of differential labdanoic acid compositions in the original resins. Another source of these molecules may be the degradation of the labdane macromolecular structure of amber due to particular conditions that prevailed in the Cantabrian deposits (see below). Other diagenetic degradation products of labdanoic acids found in both amber samples are Z- and E-19-norlabda-8(20),12-dien-15-oic acids (XXXIII and XXXIV, respectively) and bisnorlabda-8(20),12-dien-18-oic acid (XXXV) (Otto and Simoneit, 2002). It is not possible to identify

all peaks found in the polar fractions of the amber extracts due to the lack of references and possible precursors. During burial, the El Soplao and Reocín amber deposits suffered the influence of hydrothermal fluids related to the La FloridaReocín Pb–Zn mine mineralization. As a consequence, we cannot discard the possibility of alternative transformation routes leading to the unusual compounds found in the amber. These alternative transformations must be added to the lack of information about the possible chemotaxonomy of extinct Cheirolepidiaceae. Despite these uncertainties, which result in the presence of some unidentified compounds, the low maturation (cf. vitrinite reflectance data, supplementary material) and excellent preservation of the organic compounds in the amber allows us to use the components of the extract to obtain chemotaxonomic information. We suggest that the tricyclic diterpenes originally biosynthesized by the main botanical precursor of amber A were dominated by phenolic abietanes, pimarane resin acids and totarol. The bicyclic diterpenes probably contained a high proportion of labdanoic acids, in particular 13-dihydroagatholic acid (XXVIII) and communic acid (XXXVI), whose polymerization leads to the typical macromolecular structure of ambers. 3.3. Terpenoid composition of co-occurring fossil leaves and paleochemotaxonomic aspects Due to the lack of Cupressaceae representatives in all the outcrops of Las Peñosas Formation (Fig. 1) and the excellent preservation and dominance of Frenelopsis material, the comparison of the terpenoid assemblage of these plant remains with those found in the amber may help to confirm the botanical origin of the fossil resin and to understand the chemosystematics of the extinct family Cheirolepidiaceae. Since no amber associated with this family has been documented to date (Bray and Anderson, 2008; Pereira et al., 2009), the inclusion of Frenelopsis genera as a possible source of one of the amber types found in the El Soplao deposit has to be considered. Previous reports relating a possible botanical origin of ambers to Cheirolepidiaceae should be mentioned (Gomez et al., 2002b; Roghi et al., 2006). Macrofossil evidence of two potential conifer resin producers was found by Najarro et al. (2009) in the study zone: Frenelopsis (Cheirolepidiaceae family) and Mirovia (Miroviaceae). Also, the palynological record shows a contribution from the Araucariaceae family, but no meso- or macrofossils of this family have been recognized yet. As Anderson (2006) pointed out, the correlation between plant fossil evidence and co-deposited amber should be taken with caution since the major resin producer could be a minor species in the ecosystem. As we find two different potential palaeobotanical contributors for the ambers of the El Soplao deposit, all the types of plant macrofossil remains identified in the deposit were examined separately in order to establish possible chemosystematic relationships. Overall, despite the increase of aromatized derivatives such as retene, the diterpene speciation in the Frenelopsis leaves shows that all the main components are shared with the type A ambers from El Soplao (Fig. 7A). Cadalene (XXXVII) and 16,17,18-trisnorabieta-8,11,13-triene (XXIII) are among the biomarkers detected in the fossil leaves of Frenelopsis. The diagenetic processes undergone by terpenoids from the fossil leaves are consistent with those observed in sediments, as the leaves are not protected by the polymeric structure of the amber. The formation of aromatic derivatives may be governed by clay catalysis or other abiotic processes in soils and sediments (Otto et al., 2007). Also, the aromatic abietanes may be generated under aerobic conditions, consistent with the major presence of pristane and the lack of phytane (Peters et al., 2005). The presence of norabietanes is consistent with the diagenetic processes for terpenoids described by Frenkel and Heller-Kallai (1977). 14-Methyl-16,17-bisnordehydroabietane (XXVII)

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and 2,4,8-trimethylalkyltetralins (e.g. XXXI), described above, are also present in the Frenelopsis leaves. The phenolic abietane ferruginol (VIII) and its derivative 12-hydroxysimonellite (X) are key biomarkers also found in the Frenelopsis leaves. This result is consistent with the presence of ferruginol in Cenomanian Frenelopsis alata (Nguyen Tu et al., 2000a). This evidence, together with the absence of ferruginol in Mirovia leaves (Fig. 8), suggests that Frenelopsis could be one of the botanical origins for the Cantabrian amber. The terpenoid composition found in Mirovia leaves is dominated by oxidized non-specific abietane terpenoids (mainly simonellite and retene). The Frenelopsis leaves and the amber of the El Soplao deposit are largely associated with jet (black amber). The analysis of jet extract shows a composition dominated by cadalene and alkyl derivatives of naphthalene and tetralin. The identifiable terpenoids include aromatized abietanes and ferruginol. Fractionation and derivatization and GC–MS of jet extract showed the presence of ferruginol and totarol, suggesting that jet has the same botanical origin as the main type A amber in the deposit. The azulene hydrocarbon derivative guaiazulene (XXXVIII), an isomer of cadalene (XXXVII), with a strong blue color and purple-blue fluorescence, is found in low amounts in all type A ambers, Frenelopsis leaves and jet, suggesting a common origin from sesquiterpenoids synthesized by Cheirolepidiaceae. Guaiazulene is a common compound with low chemosystematic value, but the presence of this hydrocarbon in amber has not been reported to date. The significant quantity of this compound in the El Soplao samples could be at the cause for the characteristic blue-purple tinge of these ambers. Although the relationships between Cheirolepidiaceae and extant conifers are unclear (Bray and Anderson, 2008; Pereira et al., 2009), a morphological and histological correlation between Cheirolepidiaceae and Cupressaceae has been established (Daviero et al., 2001; Farjon, 2008). Moreover, Nguyen Tu et al. (2000b) have observed a resemblance between the lipid composition of Frenelopsis alata and Tetraclinis articulata, a representative of Cupressaceae. The presence of ferruginol in Frenelopsis (Nguyen Tu et al., 2000b and the present data) confirms the hypothesis of a possible relationship between Frenelopsis and the

Cupressaceae family. Moreover, the presence of 13-dihydroagatholic acid (XXVIII) in the amber and the overall biomarker assemblage show a similarity to extant Cupressus genera (see Fig. S4, supplementary material). The resemblance in the chemical composition between Frenelopsis and Cupressaceae representatives may be due to convergence, as it has been demonstrated that the physical similarities between these taxa resulted from convergence rather than phylogenetic connection (Broutin and Pons, 1975; Alvin, 1982). The evolutionary changes in the biochemistry of terpenoids since the synthesis of the parent resin of amber to the modern conifers are unknown. Consequently, we should consider that the lack or presence of certain compounds in a correlation with extant conifers is informative, and that detailed biomarker compositions of extinct conifer fossils, complemented by morphological and histological relationships, are necessary to establish a deﬁnite evolutionary relationship. Keeping this in mind, the paleobotanical considerations suggested by our data obtained on macrofossil plant samples and amber types can be summarized as follows: a. The absence of abietic and dehydroabietic acid in both types of amber samples excludes an origin from resin of the Pinaceae family. Also, the absence of plant triterpenoids and labdenoic acids eliminates a contribution from angiosperms (Anderson et al., 1992; Yamamoto et al., 2006). b. The presence of phenolic terpenoids (ferruginol, totarol and hinokiol) in the type A amber indicates the conifer families Cupressaceae and Podocarpaceae as possible biological precursors, and rejects the Araucariaceae family. The presence of callitrisic acid (XIII) reinforces the biochemical relationship between the parent resin of amber and modern Cupressaceae. The plant macrofossil record in the deposit shows that there are no representatives of Cupressaceae or Podocarpaceae among the possible resin producers (Najarro et al., 2009). The co-occurrence of key terpenoids (e.g. ferruginol), between amber A and fossil tissue of Frenelopsis suggests that this amber could be derived from Frenelopsis (Cheirolepidiaceae).

Fig. 8. GC–MS TIC trace of the total extract of Mirovia sp. leaves found in the amber deposit at El Soplao showing the identiﬁed biomarkers. Solid dots: n-alkanes (last dot: C30).

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I. 18- and 19-Norabietatrienes

II. Dehydroabietane

III. Fichtelite

IV. 18-Norabieta-7,13diene

OH V. Norabiet-13-ene

VI. Abietic acid

VII. Dehydroabietic acid

VIII. Ferruginol

OH OH

HO IX. 6,7-Dehydroferruginol

X. 12-Hydroxysimonellite

XII. Hinokiol

XI. Totarol

HO O XIII. Callitrisic acid

XIV. 18-Norferruginol

XV. Simonellite

XVI. 16,17-Bisnorsimonellite

OH OH

O XVII. Retene

XVIII. Sempervirol

XIX.Sugiol

XX. Diaromatic totarane

Fig. A1. Chemical structures cited in the text.

c. The overall terpenoid composition of the type B amber is comprised of non-speciﬁc conifer biomarkers. The absence of phenolic terpenoids and of 13-dihydroagatholic acid (XXVIII), together with the presence of major amounts of diagenetic products of pimarane-type diterpenoids, saturated and unsaturated norabietanes, and alkyltetralins point to a different biological origin. A paleobotanical source for this type of amber could not be determined on the basis of its biomarker composition. 4. Conclusions Analysis of the polar diterpenoids of Cretaceous ambers from El Soplao (Cantabria, Spain) indicates that two resin producers contributed to the amber record. The main parent resin (type A) originally contained phenolic abietanes (dominated by ferruginol), totarol, dehydroabietane and pimaric/isopimaric acids. The dominant resin acids found are 13-dihydroagatholic and bisnordehydroabietic acids, with various other alteration products and a minor quantity of callitrisic acid. This composition suggests a biochemical relation with the resin of extant Cupressaceae. The second parent resin (type B) contains pimaric/isopimaric acids as the only identiﬁable biological precursors preserved. The phenolic diterpenoids present in the

samples (type A), the lack of phyllocladane/kaurane-type terpenoids and the absence of macrofossil plant remains exclude a signiﬁcant contribution of Araucariaceae to the amber. Diagenetic products of the pimarane/abietane and labdane class terpenoids constitute the main geoterpenoids extractable from the amber of El Soplao. Insights from petrographic characterization of coal macerals provide a correlation between temperature, time and level of organic diagenesis, indicating only a moderate degree of diagenetic alteration during burial. This is consistent with the high level of preservation of the natural product diterpenoids and their direct diagenetic derivatives. The sedimentological relationships and chemotaxonomical observations suggest that one source of the amber may be the extinct Frenelopsis (Cheirolepidiaceae). Acknowledgments We thank Drs. Philippe Schaeffer, Thanh Thuy Nguyen Tu and Débora de Almeida Azevedo for their thorough and constructive comments, aiding our revision of the manuscript. We thank the staff of the Royal Botanic Garden of Madrid for permission and assistance to sample several of their living conifer species. We also thank the Consejería de Cultura, Turismo y Deportes (Gobierno de Cantabria) and in particular F.J. López Marcano

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XXI. Pimaric acid

XXIV. 16,18-Bisnorabieta-8,11,13-triene

XXIII. 16,17,18-Trisnorabieta-8,11,13-triene

XXII. Isopimaric acid

OH O

XXV. 16,17-Bisnordehydroabietic acid

XXVI. Pimar-8-en-18-oic acid

XXVII. 14-Methyl-16,17bisnordehydroabietane

HO XXVIII. 13-Dihydroagatholic acid

OH O

HO XXIX. 13-Dihydro-19-noragathic acid

XXXI. 2,5,8-Trimethyl1-butyltetralin

O XXX. Agathic acid

XXXII. 2,5,8-Trimethyl1-isopentyltetralin

OH O

OH O XXXIII. Z-19-norlabda8(20),12-dien-15-oic acid

XXXIV. E-19-norlabda8(20),12-dien-15-oic acid

XXXVII. Guaiazulene

XXXV. 14,15-Bisnorlabda-8(20),12-dien18-oic acid

HO O XXXVI. Communic acid

XXXVIII. Cadalene

Fig. A1 (continued)

(Regional Minister of the Cantabria Government) and F. Unzué (manager of El Soplao Cave and Territory) for their support and promotion of the study of the new amber deposit at El Soplao. Thanks should go to Antonino Bueno Yanes for providing Reocín amber samples. We thank the Centro de Astrobiología (CSICINTA) and Instituto de Tecnica Aerospacial ‘‘Esteban Terradas”, where all chemical analyses were performed. This study is framed in a collaborative agreement among IGME, SIEC S.A. and the Cantabria Government (Consejería de Cultura, Turismo y Deportes), and is a contribution to IGME project 491-CANOA 53.6.00.12.00: ‘‘Investigación cientíﬁca y técnica de la Cueva de El Soplao y su entorno geológico”, and DGI project CGL200801237/BTE (MICINN, Spanish Government). Appendix A See Fig. A1.

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Earth and Planetary Science Letters j o u r n a l h o m e p a g e : w w w. e l s e v i e r. c o m / l o c a t e / e p s l

Association between catastrophic paleovegetation changes during Devonian–Carboniferous boundary and the formation of giant massive sulﬁde deposits Cesar Menor-Salván a, Fernando Tornos b,⁎, David Fernández-Remolar a, Ricardo Amils a,c a b c

Centro de Astrobiología, CSIC-INTA, 28850 Torrejón de Ardoz, Madrid, Spain Instituto Geológico y Minero de España, Ríos Rosas 23, 28003 Madrid, Spain Centro de Biología Molecular “Severo Ochoa” (CSIC-UAM), Cantoblanco, 28049 Madrid, Spain

a r t i c l e

i n f o

Article history: Received 4 June 2010 Received in revised form 13 September 2010 Accepted 19 September 2010 Available online xxxx Editor: M.L. Delaney Keywords: Iberian Pyrite Belt volcanogenic massive sulﬁdes higher plant biomarkers Devonian–Carboniferous boundary

a b s t r a c t The Iberian Pyrite Belt (SW Iberia) is one of the largest sulfur anomalies in the Earth's crust. In the southern Iberian Pyrite Belt, more than 820 Mt of exhalative massive sulfides were deposited in less than one million years at the Devonian–Carboniferous boundary. The shale of the ore-bearing horizon contains biomarkers indicating major biogenic activity in a methanogenic setting, including a five-fold increase in typical vascular plant biomarkers and a significant anomaly in those probably indicating the presence of thermophilic Archaea. This contrasts with signatures in the average sedimentary rocks of the basin that indicate the sediments settled in oxic to sub-oxic environments, and that they have only minor biomarkers derived from continental paleoflora. These data show that the formation of the mineralization was not only related to major hydrothermal activity synchronous with volcanism but may also have been controlled by the input of large amounts of organic matter, mostly derived from the degradation of woodland detritus sourced in the nearby continent. This massive influx of organic matter could have accelerated extremophilic microbial activity that used short-chain hydrocarbons as electron donors for seawater sulfate reduction, resulting in concomitant massive sulfide precipitation. We propose that the giant massive sulfide deposits resulted from overlapping of geological and biological processes that occurred at the Devonian to Carboniferous transition, including: (1) continent collision during the onset of the Variscan orogeny leading to major paleogeographic changes and volcanism; (2) dramatic stress of continental ecosystems due to the combination of climatic change, volcanism, variations in the sea level and erosion on a regional scale; (3) major biomass destruction and increase of organic supply to marine environments; and, (4) generation of anoxic conditions and the thriving of sulfate reducing microorganisms. Under these conditions, massive sulfide deposits formed where venting sulfur-poor but metal-rich hydrothermal brines flowed into a hydrogen sulfide-rich anoxic water column. The data presented strongly suggest that there was a temporal and causal relationship between the Devonian– Carboniferous geotectonic, climatic and biological crises and the formation of the giant volcano-sedimentary massive sulfide deposits of the southern Iberian Pyrite Belt. © 2010 Elsevier B.V. All rights reserved.

1. Introduction The Iberian Pyrite Belt (IPB) (Fig. 1) hosts one of the largest concentrations of volcano-sedimentary massive sulﬁde deposits on Earth, comprising 45% of the giant (N100 million tons of contained ore) deposits world-wide (Leistel et al., 1998; Tornos, 2006). Furthermore, it represents one of the largest crustal sulfur anomalies, mostly in the form of pyrite in the stratabound massive sulﬁde deposits (ca. 1730 Mt), the underlying (sub-) economic stockwork or feeder zone (ca. 300 Mt) and about 3000 Mt of hydrothermally altered volcanic and siliciclastic rocks

⁎ Corresponding author. E-mail address: f.tornos@igme.es (F. Tornos).

with disseminated pyrite (Fig. 1). The ore deposits formed in a manner similar to that of present day submarine hydrothermal systems (Herzig and Hannington, 1995; Ohmoto, 1996). However, a major difference between present day systems and ancient ones is that most recent ones occur on oxic oceanic bottoms where the massive sulfides are rapidly oxidized and destroyed if they are not covered by sediments or volcanic rocks (Herzig and Hannington, 1995). Most ancient deposits occur in anoxic settings, where the massive sulfides not only accumulated but also were preserved. In fact, there was a close association between anoxic events and formation of sedimentary Zn–Pb sulfide deposits on the seafloor (Goodfellow, 1987; Goodfellow and Jonasson, 1984). The aim of this work is to study the relationships between massive sulfide formation and the geochemistry of the hosting shale, by studying the biomarkers present in the ore-bearing and barren shale.

Please cite this article as: Menor-Salván, C., et al., Association between catastrophic paleovegetation changes during Devonian–Carboniferous 108 of 138 boundary and the formation of giant massive sulﬁde deposits, Earth Planet. Sci. Lett. (2010), doi:10.1016/j.epsl.2010.09.020

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SI F

Ossa Morena Zone

S MA

Lisbon South Portuguese Zone

Madrid

Ossa-Morena Zone

I ER SP HE

Autochthonous Iberian Terrane

LO ATE LZ sj

NC Pyrite Belt

RT SO-MI MV

AZ-LF

LC 40 km

SW Portuguese Domain

Huelva

Massive sulfide deposits >100Mt

Atlantic

Faro

Ocean

>30Mt <30Mt

Post-PaleozoicCover

CulmGroup

OceanicSedimentarySequence

Intrusiverocks (mafic&felsic)

Volcano-Sedimentary Complex

OphioliticSequence

PQGroup Fig. 1. Geological map showing the location of the major deposits described in the text and the location of the San Jorge section (sj). AZ-LF: Aznalcóllar-Los Frailes; LC: Las Cruces; LO: Lousal; LZ: La Zarza; NC: Neves Corvo; MV: Masa Valverde; RT: Rio Tinto; SO-MI: Sotiel-Migollas; TH: Tharsis. Modiﬁed from Tornos (2006).

The results presented here, in combination with geologic, isotopic, and SEM data, show the critical control exerted by biogenic activity in the formation of the volcanogenic massive sulﬁde deposits in the Iberian Pyrite Belt. These deposits formed during unusual paleogeographic conditions, coeval with a paleoclimatic crisis and active hydrothermal activity, producing an explosion of microbial life and formation of giant ore deposits. 2. Geologic setting The massive sulﬁde deposits of the IPB formed synchronously with calc-alkaline volcanism during the early stages of the Variscan orogeny, when Avalonia started to amalgamate with Gondwana during late Devonian to early Carboniferous times (Silva et al., 1990). Continent–continent oblique collision led to the formation of pullapart basins on both terrains. In the exotic Avalonia terrane the onset of deformation produced tectonic disturbance of a stable continental shallow platform where mature siliciclastic sediments, shale and quartz-arenite of the PQ Group (upper Givetian to Famennian) were deposited. Formation of half-graben basins led to heterogeneous and high energy sedimentation, represented by the local deposition of

subaerial, fan delta, carbonate and gravity flow deposits during the uppermost Famennian (Moreno et al., 1996). These rocks are in part coeval with the Volcano-Sedimentary Complex, of late Devonianmiddle Visean age, that comprise large andesite–rhyolite dome complexes and basaltic lava flows and sills interbedded with abundant shale and epiclastic and chemical sedimentary rocks (Leistel et al., 1998; Tornos, 2006). The style of mineralization of the massive sulfide orebodies shows a major paleogeographic control. Massive sulfide deposits located in the northern part of the belt formed by sub-seafloor replacement of pumice- and glass-rich volcanic rocks, are unrelated to shale and are of early Visean age (Fig. 2); they are not further discussed in this paper. Those in the southern part are very different, since are hosted by shale, being exhalative in origin and formed in anoxic third order basins. Some of these deposits show abundant sedimentary structures indicating dominant precipitation on the seafloor as well as common mound-like structures made up of laminated and breccia-like pyrite– siderite rocks (Fig. 2) (Almodovar et al., 1998; Relvas et al., 2001; Tornos et al., 1998; 2008). The giant deposits of Neves Corvo, Tharsis, Masa Valverde, Aznalcóllar-Los Frailes, Las Cruces and Sotiel-Migollas, comprising in total about 820 Mt of massive sulfides, formed

Fig. 2. Relationships between geology, age, geotectonic evolution, paleogeography, and massive sulfide deposits in the Iberian Pyrite Belt. Chronostratigraphy based on the palynological data of Gonzalez (2005) and U–Pb dating of Dunning et al. (2002) ((1) in the figure) and Barrie et al. (2002) (2). The synorogenic facies of the D–C transition include the subaerial, fan delta, carbonate and gravity flow deposits of the uppermost PQ Group (Moreno et al., 1996). The column includes the location of the shale-hosted and volcanichosted massive sulfide deposits, of late Famennian and late Tournaisian age, respectively.

Please cite this article as: Menor-Salván, C., et al., Association between catastrophic paleovegetation changes during Devonian–Carboniferous boundary and the formation of giant massive sulﬁde deposits, Earth Planet. Sci. Lett. (2010), doi:10.1016/j.epsl.2010.09.020 109 of 138

311.7±1.1

foreland basin

Pennsylvanian Serpuk. Bashkir.

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318.1±1.3

Culm Group: Shale & sandstone (late Visean-late Moscovian) >2500 m

???

359.2±2.5

pull apart basin

345.3±2.1

felsic volcanic rock (Lagoa Salgada) (2) felsic volcanic rock (Las Cruces) (2) felsic volcanic rock (Riotinto) (2)

volcanic-hosted massive sulfide deposits

rhyolite (Zufre) (1) felsic volcanic rock (Aljustrel) (2)

NM Zone

rhyolite (Nerva) (1)

purple shale

shale-hosted massive sulfide deposits LN Zone

anoxia ca. 363

Famennian

stable continental platform

Transition Unit

Frasnian

375±2.6

Volcano-Sedimentary Complex: Basalt to rhyolite interbedded with shale and chemical sediments, including massive sulfides (late Famennian-early Visean) 0-1300 m

Transition Unit: subaerial, fan delta, carbonate and gravity flow deposits of the uppermost PQ Group and D-C boundary

PQ Group: shale & sandstone (late Devonian) >2000 m

385.3±2.5

Givetian

Upper Devonian

Strunian

Mississipian Tournaisian

Visean

328.3±1.6

???

Please cite this article as: Menor-Salván, C., et al., Association between catastrophic paleovegetation changes during Devonian–Carboniferous 110 of 138 boundary and the formation of giant massive sulﬁde deposits, Earth Planet. Sci. Lett. (2010), doi:10.1016/j.epsl.2010.09.020

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synchronously with, or just after the first recorded volcanic activity in the area. There are few macrofossils but systematic palinological dating of the shale hosting these orebodies (Gonzalez et al., 2006; Oliveira et al., 2003; Pereira et al., 1996 and references therein) shows that the shale-hosted massive sulfides formed during the LN biozone, the uppermost biozone of the Strunian (363–359 Ma, latermost Devonian) (Streel et al., 2006). This indicates that the shale-hosted massive sulfide deposits formed in a brief time span of less than 1 Ma and just before (if not within) the Devonian–Carboniferous (D–C) boundary, coinciding with major paleogeographic changes and the onset of volcanic activity (see Moreno et al., 1996; Silva et al., 1990) (Fig. 2). Fluid inclusion data indicate that the ore forming fluids had salinities between 3 and 12 wt.% NaCl eq. Calculated δ18O-δD values of ore forming fluids, in the range of 0 to 8‰ and −45 to 5‰ respectively, and Pb (μ = 9.8–10), Sr (87Sr/86Srinitial, 0.7021–0.7221), and Os (187Os/188Osinitial ≈ 0.69) initial values of the mineralization and related hydrothermal alteration are consistent with equilibration of fluids at low fluid/rock ratios with long lived and evolved continental crust, represented by the sedimentary rocks of the PQ Group and/or the underlying basement (Tornos, 2006). The ore forming fluids were not modified, convected and heated seawater that interacted with sediments and volcanic rocks leaching sulfur and metals, as stated for most volcanic-hosted massive sulfide deposits (e.g., Herzig and Hannington, 1995; Lydon, 1988; Ohmoto, 1996), but are interpreted to be derived from the rapid dewatering and accelerated diagenesis/low grade metamorphism of the underlying siliciclastic sequence, probably directly related to crustal thinning, magma intrusion and increase of the geothermal gradient (Tornos, 2006). Numerical modeling suggests that hydrothermal fluids that equilibrated with shale containing intermediate redox assemblages are metal-rich but sulfur-poor; these fluids travelled through the crust until they were mixed with a H2S-rich fluid, leading to the precipitation of sulfides and the formation of the massive sulfide deposits (Tornos and Heinrich, 2008). Saline fluids discharged from hydrothermal vents accumulated as brine in bathymetric depressions on the seafloor (Sato, 1972) at temperatures between 60 and 120 °C (Tornos et al., 1998). This temperature range is close to that of steady state equilibrium in brine pools (Solomon et al., 2004) and to the optimal growth conditions for thermophilic microorganisms (Stetter, 1999). Data supporting this model include the systematically low and variable δ34S (between −34 and + 21‰) values in the shale-hosted massive sulfide deposits if compared with those in the stockwork zones (−2.5 to +10‰) (Tornos, 2006; Velasco et al., 1998) and the presence in the stockwork of mineral assemblages typical of sulfurpoor environments, such as tellurides, pyrrhotite, Co–As-bearing minerals and arsenopyrite that are not found in the overlying stratabound massive sulfide deposits (Marcoux et al., 1980; Tornos et al., 1998). The sulfur isotope values of the exhalative ores are interpreted as reflecting mixing of sulfur from two contrasting sources, one of deep origin and another produced by biogenic seawater sulfate reduction under fluctuating open and closed conditions within an anoxic basin. The deep upflowing hydrothermal fluids leached metals and limited amounts of sulfur from the underlying sediments; mass balance calculations suggest that up to 50% of the sulfur was added at the site of massive sulfide formation. A biogenic source of this sulfur is consistent with an evidence of widespread biologic activity in the ore forming environment, including the presence of probable worm burrows some cm above the footwall of the massive sulfides, the formation of siderite–pyrite laminated mounds interpreted as fossil microbial mats (Tornos et al., 2008) and the presence of microfossils that may represent pyritized microbial shells. These late structures (Fig. 3) are likely of organic origin and too small to be spores.

Although this model may explain the formation of the massive sulfide deposits, it does not satisfactorily explain the mass and charge balance, as biogenic reduction of seawater sulfate requires a reducing agent. The most likely one is methane, but Tornos and Heinrich (2008) concluded that the deep hydrothermal fluids must be weakly reduced to oxidize in order to transport large amounts of metals, thus ruling out the presence of high contents of CH4. Therefore, one of the key questions related to massive sulfide formation in the IPB is the source and nature of the reductant. The drillhole PG-1 (37°36′34″N; 7°11′58″W) is located about 8 km west of the giant Tharsis deposit with about 115 Mt of massive sulfides, mostly as pyrite. This 290 m long drillhole intersected a 10 m thick pyrite-rich massive sulfide lens within a 120 m thick sequence of shale belonging to the lowermost Volcano-Sedimentary Complex that is interpreted as being a lateral equivalent of the Tharsis orebody (Tornos et al., 2008). In both cases, the ore-bearing interval coincides with a zone of major chemical changes at the basin scale (Fig. 5). The shale shows a sharp increase in the Na2O/K2O ratio, probably related to the input of detritus from the growth of the first volcanic edifices within the basin, accompanied by a dramatic drop in the Mn and Fe contents and an increase in the V/Cr ratios that are indicative of anoxic conditions. Highly reducing, even methanogenic, conditions are inferred from the high V (N300 ppm) contents, high V/Cr (N4.3), V/Ni (N4.7) ratios and low Mn (b260 ppm) contents and the absence of barite in the ore assemblage. Furthermore, the shale shows a synchronous decrease in δ34S values from the highly variable values of −21 to 3‰ of the regional shale to constant and low values of −29 to −27‰ in the ore-bearing horizon (Tornos et al., 2008). These data are consistent with ore formation taking place during a short-lived highly anoxic event that was coeval with major microbial activity and the onset of the volcanism. Shale ca. 1–2 m above the massive sulfides has V, Ni, Mn, Fe and Cr contents typical of shale deposited in sub-oxic to oxic conditions indicating that anoxic conditions were thus strictly coeval with the precipitation of the massive sulfides. The results presented here are interpreted as having a regional significance, since the geochemistry of shale is shown to be homogeneous at the basin scale (e.g., Nägler, 1990), something that is also consistent with the existence of a well defined ore forming horizon all along the southern Iberian Pyrite Belt. 3. Methods 3.1. Samples Core from drillhole PG-1 stored at the facilities of the IGME at Peñarroya (Spain) was systematically sampled by sawing half cores of the shale-rich units. About ½ kg of whole rock from 56 samples was later powdered and analyzed for major and trace elements, stable sulfur isotopes and total organic carbon (TOC). 3.2. Extraction and fractionation of the organic matter About 80 g of pulverized sample was placed in a reactor with a mixture of dichloromethane and methanol (7:1 v:v) for 8 h using a programmable hot extraction system (Büchi B-811 model). Samples were concentrated to 1 ml and desulfurized in an elemental copper powder column previously activated with 2 M HCl. After removing elemental sulfur, the solvent was removed to yield the extract for analysis. The crude extract was placed at the top of a silica column (2 ml, filled with chromatography grade silica, 0.063–0.2 mm particle diameter). Saturated and aromatic fractions were eluted with 10 ml of n-pentane:dichloromethane (4:1). A second polar fraction was eluted with 10 ml of pure dichloromethane. The separated fractions were dried under N2 and re-dissolved in 150 μl of dichloromethane.

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Fig. 3. SEM-EDS analysis of different samples obtained in the Tharsis Filón Norte massive sulfide orebody show evidences of microbial activity supporting sedimentary origin. A) Framboidal structure covered partially by a Si-rich coating with some traces of carbonaceous matter. B) Organic microfossil of unknown origin with spinose projections resembling bacterial spores or small-sized acritarchs. C) Pyritic framboid covered by a carbonaceous lamina and which mineral units are internally embedded inside an organic matrix. D) EDS spectrum obtained from the framboid shown in (C), where is compositionally dominated by sulfur and iron (pyrite) and organic carbon (no oxygen detected) supporting the occurrence of organics associated to some framboidal structures. E) Pyrite framboid associated to a spinose organic-walled microfossil (with arrow), where some micron-sized mineral units show apicular termination (see black arrows) as observed in some microbial structures casted by mucilaginous films (Gong et al., 2008).SEM-EDS analysis of ore from the Tharsis Filón Norte massive sulfides showing evidences of microbial activity. A) Framboidal structure coated swith silica (black arrows) with traces of carbonaceous matter. B) Microfossil of unknown origin with spinose projections resembling fungi as found by Jackson et al. (2009). C) Pyrite framboid covered by a carbon-rich film with the individual crystals embedded in organic matter. D) EDS spectrum of the framboid shown in (C), composed of sulfur, iron and organic carbon (no oxygen detected) supporting the occurrence of organic matter associated with some framboidal structures. E) Framboid of pyrite associated with a spinose organic-walled microfossil (arrow), where some micron-sized grains show apicular terminations (black arrows) similar to those of some microbial structures cast by mucilaginous films (Gong et al., 2008).

3.3. Gas chromatography-mass spectrometry (GC-MS) The analyses were performed on an Agilent 6850 GC coupled to an Agilent 5975 C quadrupole mass spectrometer. Separation was achieved on an HP-5MS column coated with (5%-phenyl)-methylpo-

lysiloxane (30 m × 0.25 mm, 0.25 μm ﬁlm). The operating conditions were as follows: 8 psi carrier pressure, initial temperature held at 40 °C (1.5 min), increased from 40 to 150 °C at a rate of 15 °C/min, held 2 min, increased from 150 to 255 °C at a rate of 5 °C/min, held for 20 min and a ﬁnal increase to 300 °C at a rate of 5 °C/min. The sample

Please cite this article as: Menor-Salván, C., et al., Association between catastrophic paleovegetation changes during Devonian–Carboniferous 112 of 138 boundary and the formation of giant massive sulﬁde deposits, Earth Planet. Sci. Lett. (2010), doi:10.1016/j.epsl.2010.09.020

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was injected in splitless mode with the injector temperature at 290 °C. The mass spectrometer was operated in electron impact mode (EI) at 70 eV ionization energy, scanned from 40 to 700 Da and operated in SIM mode at m/z 85, 183, 219 and 197. Temperature of the source was 230 °C and quadrupole temperature was 150 °C. Data were acquired and processed with Chemstation software. Individual compounds were identiﬁed by comparing their mass spectra with those of standards and with published data. 4. Organic geochemistry of the ore-bearing sequence Biomarker analysis of rock samples from the PG-1 drillcore was performed to elucidate the biological sources of the organic matter in the basin. Shale from the ore-bearing horizon contains 0.045 ± 0.005% of mean extractable organic matter which is composed mainly of saturated, linear and branched hydrocarbons, unspeciﬁc aromatic compounds and bacterial and plant biomarkers (Fig. 4). A clear quantitative increase in the biomarkers produced by plant degradation was observed in the ore-related samples when compared with barren rocks below and above the orebody (Fig. 5).

et al., 2005; van Aarssen et al., 1996, 2000). The Higher Plant Index was calculated using the equation: −1

HPI = ð½retene + ½cadalene + ½ip iHMN Þ ð½TeMN Þ

Where TeMN is the bacterially derived 1,3,6,7-tetramethylnaphthalene, used to correct the higher plant input with bacterial inputs at the same maturation grade. The quantification was performed using the mass fragments at m/z 219 for retene, m/z 183 for cadalene, m/z 197 for ip-iHMN and m/z 169 for TeMN respectively. A five-fold increase in the Higher Plant Index (HPI) was observed in the ore-bearing horizon, in the same trend with the individual higher plant biomarkers (Fig. 5). The HPI increase suggests a higher input of vascular plants to the sediments. This major change is accompanied by an increase in the 2-methylretene/9-methylphenanthrene ratio; 2-methylretene has been interpreted as a gymnosperm marker through diagenesis of abietane and phyllocladane class diterpenes. To assess the possible effect of differential maturation in the orebody, the values of the biomarker were corrected using the ratio with the unspecific maturation product 9-methylphenanthrene (Bastow et al., 2001).

4.1. Higher Plant Index 4.2. Branched alkanes with quaternary carbon The three aromatic biomarkers selected for the study of variation in higher plant input were cadalene (4-isopropyl-1,6-dimethylnaphthalene), retene (1-methyl-7-isopropylphenanthrene) and ipiHMN (6-isopropyl-1-isohexyl-2-methylnaphthalene). These three compounds have natural precursors abundant in extant plants (Peters

Relative Abundance

Cadalene

4.3. Paleovegetation proxies

C19 Pr

Relative Abundance

Shale samples from the ore-bearing horizon are signiﬁcantly enriched in 5,5-diethyl-alcane hydrocarbons including 5.5-diethyl pentadecane and 5,5-diethyl-tridecane (Fig. 6). The biosynthetic origin of quaternary diethyl alkanes has been demonstrated, but the source organisms are unclear due to the scarcity of these structures in modern environments. The presence of branched quaternary carbons in modern deep sea hydrothermal waters suggests that they may be biomarkers for non-photosynthetic reducing thermophilic prokaryotes or Archea (Kenig, 2003, 2005). Also, these molecules may be indicators of the response of source organisms to environmental changes, as euxinic to oxygenated water columns lack quaternary carbon branched alkanes, indicating low oxygen concentrations (Kenig, 2003).

Time Fig. 4. GC-MS trace in SIM mode at m/z 183 and 85 of the aliphatic/aromatic fraction. A: sample from level 84 m, showing a high relative abundance of cadalene with respect to n-alkanes. B: sample at level 26.15 m, above the mineralization, at the stabilization zone of the higher plant input. The abundance of cadalene relative to n-alkanes is lower above and below the orebody. C19: n-nonadecane; Pr: pristane; and Ph: phytane.

The analysis of organic extracts from samples obtained at the bottom of and below the mineralized horizon gave relatively low HPP (Higher Plant Parameter) values. This parameter is expressed as the abundance of retene relative to the sum of retene and cadalene, measured at m/z 219 (retene) and m/z 183 (cadalene). HPP is used as a tracer of the evolution or changes of paleovegetation, as the increase of gymnosperm input leads to an increase of HPP (Fleck, 2002; van Aarssen et al., 2000). This inference is based in the origin of cadalene, a diagenetic product of cadinene, the common product of sesquiterpenoid biosynthetic pathway and considered a generic vascular plant biomarker. Retene is the final diagenetic product of the abietane-class diterpenes, characteristic products of gymnosperm secondary metabolism. Hence, the reduced level of HPP in the orebody reflects a decrease in gymnosperm contribution to the terrigenous input of organic matter. These reduced levels could be related to severe alteration of ecosystems at the beginning of the mineralization period, followed by a recuperation of the gymnosperm dominance reflected as an increase in HPP values. The decrease of HPP in the initial stages of black shale deposition and orebody formation is consistent with the increase in cadalene in the ore-bearing horizon (Fig. 5). The ratio cadalene/n-nonadecane is used to correct the variation of cadalene in relation to variations in total extractable organic matter and errors in sample manipulations, as n-nonadecane levels relative to the total peaks found do not show

-20

Thrust fault

S (o/oo)

Organic Carbon (%)

5-5' diethylpentadecane (%)

Higher Planr Parameter

Higher Plant Index

Ratio 2MR/9MP

2-M-Retene (%)

Ratio CD/C19

Cadalene (%)

Lithology

Depth (meters)

C. Menor-Salván et al. / Earth and Planetary Science Letters xxx (2010) xxx–xxx

No record

-40 -60 -80 -100 -120 -140

-180 -200 -220

Lithology Felsic volcanic rocks Pyrite-richshale (ore-bearing horizon)

Volcano-Sedimentary Complex

-160

-240

Gray shale

-260

Shale & quartz-arenite Phyllite-Quartzite Group

-280

Fig. 5. Composite stratigraphic column of the San Jorge section showing variations of some key geochemical parameters with depth. CD/C19: ratio cadalene/n-nonadecane; 2-M-Retene: 2-methylretene; and 2MR/9MF: ratio 2-methylretene/9-methylphenanthrene.

variation along the drillcore studied. This ratio could also be indicative of the higher plant (terrigenous)/algal (marine) ratio and show the same trend in vertical variation as the terrigenous/aquatic ratio (TAR), calculated using the distribution of odd n-alkanes (Peters et al., 2005). The increase of TAR in the orebody must be taken with caution, as the n-alkane ratios are sensitive to secondary, post-depositional processes and its usage is normally restricted to young sediments. Despite these qualiﬁcations, the variation of TAR with the same trend as the other indicators is consistent with the increase of terrigenous material input to the basin. In the later stages of the mineralization period there is a second major and similar biogeochemical discontinuity. Samples above the pyritic shale-gray shale contact at ca. 55 m. in drillhole PG-1 (Fig. 5) are also characterized by an abrupt increase in HPP and dehydroabietic acid, which is a diagenetic product of abietic acid also indicative of gymnosperms (Otto and Simoneit, 2001). Given that the dehydroabietic acid and dehydroabietanes enrichments correlate with higher HPP values, and the lack of dehydroabietanes in the shale

below ca. 55 m, it can be deduced that the continental ecosystem was dominated by new species of gymnosperm-like plants during the later stages of ore formation (Fleck, 2002). 4.4. Plant biomass input sourced from terrestrial ecosystem destruction To determine the continental biomass that was consumed by microbial activity, we considered a simple biogeochemical reaction involving a six carbon molecule of glucose that is the molecular base for cellulose. This mechanism results from activation of the sulfate ion by adenosine triphosphate (ATP), which is catalysed by the enzymes ATP sulfurylase and APS reductase. In dissimilatory sulfate reduction mediated by microbes, the electrons are transferred by a hydrogenase to the low potential cytochrome c3 that releases electrons to a cytochrome complex, which is the ﬁnal stage of reduction to sulﬁde. Glucose was selected as the potential electron donor given that some hyperthermophilic and sulfate-reducing archaea can reduce sulfate through incomplete oxidation of glucose to produce acetate and CO2

Please cite this article as: Menor-Salván, C., et al., Association between catastrophic paleovegetation changes during Devonian–Carboniferous 114 of 138 boundary and the formation of giant massive sulﬁde deposits, Earth Planet. Sci. Lett. (2010), doi:10.1016/j.epsl.2010.09.020

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which can be combined into the main reaction to produce pyrite synchronous with the release of CO2 57

2þ

C6 H12 O6 þ 3•SO4 þ 1:5•Fe þ 7:5•H2 O 5,5-diethylpentadecane

Intensity

þ 3•H þ 0:75•O2 →1:5•FeS2 þ 6•CO2

The clearest evidence for this process is provided by the siderite– pyrite laminated ores and related breccias that are interpreted as fossil biogenic mounds (Tornos et al., 2008). The formation of siderite was favored by an increase in the pCO2, generated as a by-product during the microbial reduction of seawater sulfate to H2S and concomitant oxidation of organic matter to CO2 due to the simpliﬁed reactions

127 239 211

CH3 COOH þ SO4 þ 2•H ⬄H2 S þ CO2 þ 2•H2 O

Intensity

85 5,5-diethyltridecane

127

211 183

100

200

300

m/z Fig. 6. Mass spectrum and structure of 5,5-diethylpentadecane (A), corresponding to the peak observed at 18.18 min in the fractionated extracts, and 5,5-diethyltritecane, observed at 14.3 min in the fractionated extracts (B). Identiﬁcation based on data published in Kenig (2005). These compounds are the major branched alkanes with quaternary carbon (BAQCs) observed in samples.

(Labes and Schönheit, 2001). However, the acetate could be further oxidized by other sulfate reducing bacteria to CO2 as has been observed in different microbial communities (Monetti and Scranton, 1992). Having these pathways in mind the reactions to produce pyrite mediated by microbes are the following: ¼

−

C6 H12 O6 ðglucoseÞ þ SO4 →2•CH3 COO ðacetateÞ þ H2 S þ 2•CO2 þ 2•H2 O

−

CH3 COO ðacetateÞ þ SO4 þ 3•H →H2 S þ 2•CO2 þ 2•H2 O

2þ

H2 S þ 0:5Fe

þ 0:25O2 →H þ 0:5•FeS2 ðpyriteÞ þ 0:5•H2 O

ð4Þ

ð1Þ

ð2Þ

ð3Þ

ð5Þ

with the produced H2S reacting with metals and precipitating the sulfides. There is no major increase in the total organic carbon in the orebearing horizon when compared with the adjacent rocks (Fig. 5). This may indicate that amount of formed pyrite was limited by the availability of organic matter required as an electron donor in the reaction (4). Assuming that all the sulfide orebody is genetically associated with the destruction of organic matter, it is possible to calculate a maximum amount for the organic carbon that was involved in the reaction. Total original reserves of the Tharsis deposit were about 115 Mt of almost massive pyrite, of which roughly 105 Mt corresponds to sulfides and 10 Mt to siderite. This implies the presence of 56 Mt of sulfur (around 5.6×1013 g S) contained in the pyrite, which is the amount that has been used to estimate the total organic carbon consumed by sulfate reduction. Using Eq. (4), the total moles of carbon degraded by microbes is half of the total moles of sulfur, which results in 8.75 × 1011 C moles and around 1.05 × 1013C grams. Assuming that carbon in siderite is derived from the CO2 generated in the Eq. (4), direct estimations provide around 1012 C grams. Adding both sources of carbon, siderite and pyrite, we estimate that destruction of the Late Devonian ecosystems provided up to 1.15 × 1013 C grams to the thermophilic communities in less than one million of years. Some spectacular findings in early Carboniferous deposits (FalconLang, 2004) provide information that allows us to make a rough estimation of the extent of the terrestrial ecosystems that were destroyed by the geotectonic processes required to produce the deposits in the Pyrite Belt. Falcon-Lang's paper gives data concerning the size, morphology and density of trees and forests in the earliest Carboniferous ecosystems. One important parameter is the mean tree diameter that is used to calculate the total stored carbon per tree by using different equations (Aparecida Vieira et al., 2009; Chave et al., 2005). The Mississippian woodlands described by Falcon-Lang (2004) had a trunk diameter averaging 6.26 cm and a tree density of around 1.8 x 106 trees per km2. The application of the equation of Chave et al. (2005) allows us to calculate the number of trees on the basis of the total carbon consumed, as inferred from the size of the massive sulfide deposits. In this equation (LnB = −2.43 + lnD) a regression between biomass B (Kg) and tree diameter D (cm) implies up to 5.75 × 109 trees, which corresponds to an area of 3200 km2 based on the average tree density calculated by Falcon-Lang (2004).

Fig. 7. Cartoon showing the development and destruction of the late Devonian terrestrial ecosystems during the onset of the Variscan orogeny in the area. a, Diagram showing Famennian (late Devonian) ecosystems dominated by progymnosperms, lycopsids, equiseta and ferns (see Algeo et al., 2001) with different water requirements from wet to drier conditions as in some modern ecosystems. The distribution of vegetal communities would have been arranged according to the regional geomorphology that controlled the topology and depth of the phreatic level in the substrate. At this stage we believe a regular phreatic level was likely, which is consistent with stable terrestrial communities, but there were also shallow inter-deltaic marine sedimentary areas producing carbonates (Moreno et al., 1996). b, Diagram showing dramatic changes in landscape affecting the terrestrial biota by the first stages of the Variscan orogeny, which induced compartmentation by formation of pull-apart basins and active volcanism. Combination of both processes induced several mechanisms such as (1) production of acid rain and dark effect, (2) disturbance of the phreatic level and regional climate by uplift and deepening, and, (3) rapid increase in erosion rates. Together, these actions worked to increase the supply of organic matter to the ocean. The gray tones in the massive sulfides include both the carbonate ore (light) interpreted as precipitated as biogenic mounds and the fine grained ore (dark), probably formed by direct fallout of sulfides from the chemocline.

C. Menor-Salván et al. / Earth and Planetary Science Letters xxx (2010) xxx–xxx

5. Interpretation The data presented here indicate that the massive sulﬁde deposits of the Iberian Pyrite Belt formed synchronously with major changes in

the biota and a major input of tree detritus from the nearby continent into the ocean. In detail, they formed in third order basins where hot, mildly reduced, metal-rich but H2S-poor hydrothermal ﬂuids accumulated and dissimilatory biogenic reduction of seawater sulfate

Please cite this article as: Menor-Salván, C., et al., Association between catastrophic paleovegetation changes during Devonian–Carboniferous 116 of 138 boundary and the formation of giant massive sulﬁde deposits, Earth Planet. Sci. Lett. (2010), doi:10.1016/j.epsl.2010.09.020

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became widespread (Goodfellow, 1987). Supplied organic matter acted as the critical electron donor needed for biogenic sulfate reduction and ore formation. Thermophilic archaea, such as Archeoglobus spp., reduced seawater sulfate to H2S, thereby providing an almost unlimited supply of reduced sulfur to fix hydrothermal metals and form the massive sulfide deposits. Thus, input of large amounts of biomass into anoxic bottoms where submarine hydrothermal activity had already accumulated metal-bearing but sulfur-poor hot fluids favored microbial life explosion and highly effective formation of the giant massive sulfide orebodies. The recorded change can only be related to a catastrophic event and with effects similar to those recorded at the KT boundary, where there is widespread evidence of major devastation and mass plant extinction (Johnson et al., 2000; Wilf and Johnson, 2004). The absence of retene-like biomarkers in the studied shale suggests that regionally extensive fires were not the main cause of forest destruction near the IPB. The ore-bearing horizon show only a slight increase in benz[a] anthracene, a possible biomarker for burning processes, but its presence alone is not conclusive. The timing of hydrothermal activity and biomass destruction coincides with the Devonian–Carboniferous (D–C) boundary, one of the dozen major biota extinction events. The D–C boundary is related to the disappearance of about 21% of marine genera and the vanishing of sessile benthos and nektobenthic organisms, mainly most ammonoids, but also foraminifera, conodonts, brachiopods, corals and others (Caplan and Bustin, 1999; McLaren and Goodfellow, 1990; Sepkoski, 1996). This mass extinction was accompanied by an event of widespread anoxia (Hangenberg event) that has been tracked mainly at low paleolatitudes. Consistent with the Hangenberg event, our data shows an increase in branched hydrocarbons with quaternary carbon (BAQC). This increase could be the consequence of flourishing of organisms implicated in the sulfur cycle due to anoxia and increase of reduced carbon input, as BAQCs are considered indicators of environmental change to anoxic conditions (Bai et al., 2006; Kenig, 2003). Alternatives have been proposed for the D–C global extinction (Caplan and Bustin, 1999), including major volcanism, global cooling and warming, oceanic overturns, eutrophication, eustatic changes, or their combinations. Detailed isotopic studies show that climate warming and a rise in sea level due to melting of ice sheets followed an epoch of glaciation during the late Devonian (Kaiser et al., 2006). In such conditions, water circulation should be inhibited, especially in subequatorial zones (Eastoe and Gustin, 1996), favoring eutrophication. Initiation of a short-lived regional anoxia event could be triggered by massive biomass extinction, favored by major input of Corg from land and an increase of δ13C values as has been recorded by Kaiser et al. (2006). In fact, the observed increase in the global δ34S values of seawater sulfate at the D–C boundary (Caplan and Bustin, 1999; Claypool et al., 1980) can be attributed to high rates of biogenic reduction and precipitation of pyrite during this anoxic period. Furthermore, McLaren and Goodfellow (1990) have claimed that most of these catastrophic events are related to meteoritic impacts, but no geologic nor geochemical evidence of such an event has been found in the Iberian Pyrite Belt. Our hypothesis is that volcanism was accompanied by dramatic climatic changes at a regional scale. The onset of extensive subaerial volcanic activity was accompanied by an increase in the geothermal gradient and the formation of large clouds of volcanic ash and dark winter as well as acid rain that caused the devastation of large forests and further extensive erosion. This should produce a dramatic decrease of continental biomass, ocean eutrophication, formation of anoxic bottoms and onset of favorable conditions for massive biogenic reduction of seawater sulfate. Alternatively, rapid topographic changes related to basin configuration during oblique collision, with depression–submersion or uplift of large areas, may also have produced severe deforestation by the reconfiguration of the regional hydrology and fluvial networks. The destruction of the forests may

also have resulted in an over-flux of atmospheric CO2, which would not have been readily buffered due to widespread biomass loss in the ocean. Preserved volcanogenic massive sulfide deposits of late Devonian– early Carboniferous age are mainly restricted to the Iberian Pyrite Belt and the D–C transition is not an epoch of relevant global volcanic activity. All these processes are likely interconnected to the global changes that led to the mass extinction of plants and animals at the D– C boundary. However, to what extent this forest devastation had global significance is outside the scope of this paper. Despite being an attractive option, we cannot be certain that the geographically restricted volcanism and hydrothermal activity had such a dramatic impact on the global climate during the D–C boundary. More likely, volcanism and hydrothermal activity are a consequence of the major geotectonic changes that took place at the D–C transition in the region of the Iberian Pyrite Belt (Fig. 7).

6. Conclusions The geochemistry of shale hosting the giant massive sulfide deposits of the southern Iberian Pyrite Belt shows that there is a close relationship between the presence of biomarkers indicative of vascular plants, and the ore deposits. The formation of the massive sulfide deposits is believed to be the result of the microbial activity driven by the degradation of plant organic matter sourced on the adjacent continent. The appearance of large amounts of tree detritus coincides in time with the D–C boundary and the onset of volcanism in the area, suggesting a relationship between catastrophic events, major geotectonic changes and formation of a large ore province. The formation of the giant massive sulfide deposits of the Iberian Pyrite Belt was due to the unusual combination of different crustal processes that took place at or near the Devonian–Carboniferous boundary, with superposition of likely independent, but probably intimately interconnected macroto microscale geologic processes that occurred in a time span of less than 1 Ma. They include early continent collision, formation of shallow marine basins, volcanism and hydrothermal activity during the early part of the Variscan orogeny, dewatering of sedimentary basins, formation of anoxic bottoms, extensive biomass destruction and flourishing of extremophile biogenic activity. The availability of large amounts of biogenically derived sulfur was the ultimate key for ore formation in the Iberian Pyrite Belt. Without that process, controlled itself by the devastation of large continental forests, metals contained in the hydrothermal fluids would have been lost from the basin, preventing the formation of these giant deposits. Thus, our data strongly suggest that the massive sulfide deposits of the Iberian Pyrite Belt are directly related a major tectonic, climatic and biological crisis.

Acknowledgements This work was supported by internal projects of the CAB and the IGME and the projects DGI-FEDER CGL2006-0378, ESP2006-556 and AYA2009-11681 of the Research and Innovation Ministry of Spain. We thank A. H. Knoll, H. Ohmoto, C. Conde, W. D. Goodfellow, C. Heinrich, F. Velasco and, especially, N.C. White their suggestions and comments. We extend our sincere thanks to the editor, M.L. Delaney, and two anonymous referees for their suggestions, comments and help during the editing process.

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Acta 65, 3505–3527. Pereira, Z., Saez, R., Pons, J.M., Oliveira, J.T., Moreno, C., 1996. Edad devónica (Struniense) de las mineralizaciones de Aznalcóllar (Faja Pirítica Ibérica) en base a palinología. Geogaceta 20, 1609–1612. Peters, K.E., Walters, C.C., Moldowan, J.M., 2005. The Biomarker Guide, second edition. Volume 2: biomarkers and isotopes in petroleum exploration and Earth history. Cambridge University Press, Cambridge. Relvas, J.M.R.S., Tassinari, C.C.G., Munha, J., Barriga, F.J.A.S., 2001. Multiple sources for ore forming fluids in the Neves Corvo VHMS deposit of the Iberian Pyrite Belt (Portugal): Strontium, Neodymium and Lead isotope evidence. Mineralium Deposita 36, 416–427. Sato, T., 1972. Behaviours of ore forming solutions in seawater. Min. Geol. 22, 31–42. Sepkoski, J.J., 1996. Patterns of Phanerozoic extinction: a perspective from global data bases. In: Walliser, O.H. (Ed.), Global Events and Event Stratigraphy in the Phanerozoic. Springer, Berlin, pp. 35–51. Silva, J.B., Oliveira, J.T., Ribeiro, A., 1990. Structural outline of the South Portuguese Zone. In: Dallmeyer, R.D., Martinez García, E. (Eds.), PreMesozoic Geology of Iberia. Springer, Heidelberg, pp. 348–362. Solomon, M., Tornos, F., Large, R.R., Badham, J.N.P., Both, R.A., Zaw, K., 2004. Zn–Pb–Cu volcanic-hosted massive sulphide deposits: criteria for distinguishing brine pooltype from black smoker-type sulphide deposition. Ore Geol. Rev. 25, 259–284. Stetter, K.O., 1999. Extremophiles and their adaptation to hot environments. FERS Lett. 452, 22–25. Streel, M., Brice, D., Mistiaen, B., 2006. Strunian. Geol. Belg. 9, 105–109. Tornos, F., 2006. Environment of formation and styles of volcanogenic massive sulfides: the Iberian Pyrite Belt. Ore Geol. Rev. 28, 259–307. Tornos, F., Gonzalez Clavijo, E., Spiro, B.F., 1998. The Filón Norte orebody (Tharsis, Iberian Pyrite Belt): a proximal low-temperature shale-hosted massive sulphide in a thin-skinned tectonic belt. Miner. Deposita 33, 150–169. Tornos, F., Heinrich, C.A., 2008. Shale basins, sulfur-deficient ore brines, and the formation of exhalative base metal deposits. Chem. Geol. 247, 195–207. Tornos, F., Solomon, M., Conde, C., Spiro, B.F., 2008. Formation of the Tharsis massive sulfide deposit, Iberian Pyrite Belt: geological, lithogeochemical, and stable isotope evidence for deposition in a brine pool. Econ. Geol. 103, 185–214. Van Aarssen, B.G.K., Alexander, R., Kagi, R.I., 1996. The origin of Barrow Sub-basin crude oils: a geochemical correlation using land plant biomarkers. APPEA J. 36, 465–476. Van Aarssen, B.G.K., Alexander, R., Kagi, R.I., 2000. Higher plant biomarkers reflect paleovegetation changes during Jurassic times. Geochim. Cosmochim. Acta 64, 1417–1424. Velasco, F., Sanchez España, J., Boyce, A., Fallick, A.E., Saez, R., Almodovar, G.R., 1998. A new sulphur isotopic study of some Iberian Pyrite Belt deposits: evidence of a textural control on some sulphur isotope compositions. Miner. Deposita 34, 1–18. Wilf, P., Johnson, K.R., 2004. Land plant extinction at the end of the Cretaceous: a quantitative analysis of the North Dakota megafloral record. Paleobiology 30, 347–368.

Please cite this article as: Menor-Salván, C., et al., Association between catastrophic paleovegetation changes during Devonian–Carboniferous 118 of 138 boundary and the formation of giant massive sulﬁde deposits, Earth Planet. Sci. Lett. (2010), doi:10.1016/j.epsl.2010.09.020

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Review of the El Soplao Amber Outcrop, Early Cretaceous of Cantabria, Spain María NAJARRO1, *, Enrique PEÑALVER1, Ricardo PÉREZ-DE LA FUENTE2, Jaime ORTEGA-BLANCO2, Cesar MENOR-SALVÁN3, Eduardo BARRÓN1, Carmen SORIANO4, Idoia ROSALES1, Rafael LÓPEZ DEL VALLE5, Francisco VELASCO6, Fernando TORNOS1, Véronique DAVIERO-GOMEZ7, Bernard GOMEZ7 and Xavier DELCLÒS2 1 Instituto Geológico y Minero de España, E-28003, Madrid, Spain 2 Departament d’Estratigrafia, Paleontologia i Geociències Marines, Facultat de Geologia, Universitat de Barcelona, E-08071, Barcelona, Spain 3 Centro de Astrobiología (CSIC-INTA), E-28850, Torrejón de Ardoz, Spain 4 European Synchrotron Radiation Facilities, F-38000 Grenoble, France/Géosciences Rennes - Université de Rennes 1, F-35042 Rennes, France 5 El Soplao S.L., Prao El Collao, E-39553, Cantabria, Spain 6 Universidad del País Vasco, Dpto. Mineralogíay Petrología, E-48080 Bilbao, Spain 7 CNRS-UMR 5125 PEPS, Université Lyon 1, Géode, F-69622 Villeurbanne, France

Abstract: El Soplao outcrop, an Early Cretaceous amber deposit recently discovered in northern Spain (Cantabria), has been shown to be the largest site of amber with arthropod inclusions that has been found in Spain so far. Relevant data provided herein for biogeochemistry of the amber, palynology, taphonomy and arthropod bioinclusions complement those previously published. This set of data suggests at least two botanical sources for the amber of El Soplao deposit. The first (type A amber) strongly supports a source related to Cheirolepidiaceae, and the second (type B amber) shows non-specific conifer biomarkers. Comparison of molecular composition of type A amber with Frenelopsis leaves (Cheirolepidiaceae) strongly suggests a biochemical affinity and a common botanical origin. A preliminary palynological study indicates a regional high taxonomical diversity, mainly of pteridophyte spores and gymnosperm pollen grains. According to the preliminary palynological data, the region was inhabited by conifer forests adapted to a dry season under a subtropical climate. The abundant charcoalified wood associated with the amber in the same beds is evidence of paleofires that most likely promoted both the resin production and an intensive erosion of the litter, and subsequent great accumulation of amber plus plant cuticles. In addition, for the first time in the fossil record, charcoalified plant fibers as bioinclusions in amber are reported. Other relevant taphonomic data are the exceptional presence of serpulids and bryozoans on the surfaces of some amber pieces indicating both a long exposure on marine or brackish-water and a mixed assemblage of amber. Lastly, new findings of insect bioinclusions, some of them uncommon in the fossil record or showing remarkable adaptations, are reported. In conclusion, a documented scenario for the origin of the El Soplao amber outcrop is provided. Key words: fossil resin, chemotaxonomy, paleobotany, charcoal, arthropod bioinclusions, taphonomy, Early Albian

1 Introduction Amber preserves delicate organic fossils, including microorganisms, cells and tissues, but the most abundant record is constituted by insects, sometimes showing interactions between them, such as mating interaction, commensalism and parasitism (e.g. Grimaldi, 1996; Martínez-Delclòs et al., 2004; Grimaldi and Engel, 2005). Early Cretaceous amber is remarkably important, because it was during this period that * Corresponding author. E-mail: m.najarro@igme.es

there occurred explosive radiations of the flowering plants and many modern families of insects (Grimaldi and Engel, 2005). During the last decade several new deposits of Cretaceous amber have been discovered in Spain and France (e.g. Delclòs et al., 2007; Perrichot and Néraudeau, 2009), and researches on various aspects of Cretaceous amber have increased greatly. However, most of the researches made concern the taxonomical descriptions of arthropods and microorganisms embedded in amber. Only a few studies have been conducted on the origin of both the amber and the deposits that contain it. One of the most

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important aspects concerning Cretaceous amber is the plant source, if there was only one, but this aspect has been shown to be very difficult to ascertain with confidence. The first detailed description of Spanish amber was published by Casal (1762). Approximately 150 years later Boscá (1910) first indicated the possible presence of insects preserved in Spanish amber. After the discovery of two remarkable amber deposits rich in bioinclusions, Peñacerrada I in the north and San Just in the north-east of Spain, a third major amber deposit was discovered in 2008 in Cantabria, named El Soplao. Thus it was not included in the most recent review of Spanish amber outcrops by Delclòs et al. (2007). Najarro et al. (2009) described in detail the regional geology, the features of the amber pieces, including infrared spectroscopy analyses, and a first attempt of the bioinclusions and the plant cuticles associated with the amber. Subsequent papers deal with biogeochemistry of amber and descriptions of new arthropod taxa preserved as bioinclusions. Thus, Menor-Salván et al. (2009a, 2009b) investigated biomarkers from amber pieces and Frenelopsis leaves from the El Soplao outcrop, concluding that they share the same origin; in addition, they identified the chemical compound form that produces an intensive fluorescent blue glow when this amber is under normal sunlight. More recently, Menor-Salván et al. (2010) report some paleochemotaxonomical aspects of the biological diterpenes preserved in El Soplao amber. On the other hand, Pérez-de la Fuente et al. (2010) and Nel et al. (2010) described the new insect taxa Cantabroraphidia marcanoi (Raphidioptera: Mesoraphidiidae) and Tethysthrips hispanicus (Thysanoptera: Thripidae), respectively. Lastly, Ortega-Blanco et al. (2010a) report the fauna of false fairy wasps (Mymarommatoidea) that has been recorded in Spanish Cretaceous amber, including a paratype specimen from El Soplao amber. We report relevant new data about El Soplao amber deposit, such as new detail on stratigraphy of the amber-bearing levels, new geochemical and biogeochemical information of the amber and fossil leaves, palynological data, the presence of a great abundance of charcoal associated with the amber, the presence of marine invertebrates on the surface of some amber pieces and new discoveries about the bioinclusions. These new contributions permit us to portray a more documented scenario for the origin of the El Soplao amber outcrop.

2 Geology The El Soplao outcrop belongs to the Cretaceous succession of the north-western margin of the Basque-Cantabrian Basin. During the Cretaceous, this basin was affected by extensional tectonics, and perhaps strike-slip, associated with the opening of the North Atlantic Ocean and the Bay of Biscay (e.g. Le Pichon and Sibuet, 1971; Rat, 1988; García-Mondéjar et al., 1996; Soto et al., 2007). During the Late Jurassic–Early Cretaceous a major rift phase developed that led to the formation of several narrow sub-basin controlled by E–W, NW–SE and SW–NW trending faults, in which variable thicknesses of continental to marine sediments accumulated (García-Mondéjar et al., 1996; Soto et al., 2007). The

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El Soplao area lies in the North Cantabrian sub-basin located immediately to the north of the Cabuérniga Ridge (Wilmsen, 2005), which is an E–W late-Variscan structure reactivated first as a paleo-high bounded by normal faults during the Early Cretaceous, and later as a reversal fault during widespread Tertiary (Pyrenean) compression (Fig. 1). The Lower Cretaceous (Aptian–Albian) deposits of the El Soplao area lie unconformably on lower Triassic (Buntsandstein facies) basement. They constitute a relatively thin (~300 m) sedimentary wedge weakly deformed and affected only by gently folding. The El Soplao amber outcrop is located in the southern flank of the Bielva syncline (Fig. 1), where the average strike of the succession is E– W and the dip about 40º N. Thermal maturity of the area based on vitrinite reflectance values ( % Ro) performed on vitrinite macerals from plant fragments have yield relatively low reflectance values that range between 0.50% and 0.61% (mean 0.56%) (Menor-Salván et al., 2010). These values suggest that the organic matter associated with the amber is only early mature, and the estimated temperatures from these indices of thermal maturation of organic matter are in the range of 60–70ºC (e.g. Sweeney & Burnham, 1989), which are the burial maximum temperatures suffered by the amber deposit. These low maturity levels may be responsible for the good conservation of the molecular composition of the amber and its biological inclusions. A synthesis of the stratigraphy of the Soplao area is represented in Fig. 2 (after Najarro et al., 2009, 2010). The amber-bearing deposit of El Soplao is included within the Las Peñosas Formation (Fig. 2), which is a Lower Albian unit (~112–110 Ma) of continental to transitional marine siliciclastic deposits interbedded in a succession of shallow marine, rudist and coral carbonate platform deposits. General sedimentary descriptions, depositional environments and fossil content of the Las Peñosas Formation have been already discussed in Najarro et al. (2009). In the El Soplao outcrop, the amber-rich beds occur in the lower-middle part of the Las Peñosas Formation. There, this unit rests disconformably above a thin (1–2 m) bed of continental red clay with root traces, deposited at the top of shallow marine limestones of the Reocín Formation (Fig. 2). The El Soplao amber deposit is characterized by dark, carbonaceous, pyritiferous shales with subordinated siltstones and sandstones laminae and crosslaminated centimetric sandstone layers, forming wavy and lenticular bedding. They contain remarkable accumulations of plant remains and amber pieces of different sizes and forms, as well as some remains and small shells of marine gastropods and bivalves. The principal amber-bearing shale bed of the El Soplao outcrop forms a lenticular body with maximum thickness of 1.5– 2.0 m and a width of at least 10 m in N–S cross-section. In the strike of the bed, the amber-bearing shale extends more than 75 m. The base of this len-shaped shale bed is erosional and truncates highly bioturbated, medium to coarse-grained sandstones with cross-bedding. At its top, the amber-bearing shale bed is overlain by heterolithic carbonaceous sandstones and muddy siltstones to sandstones with flaser bedding, which are also relatively rich in amber pieces. All of these deposits accumulated in a proximal estuarine bay system with small bayhead deltas (Najarro et al.,

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Fig. 1. Geological map of the study area showing the location of the El Soplao amber outcrop. Modified from Najarro et al. (2009).

2009) and represent the transgressive inundation of a continental fluvial plain and an incised valley fill. Based on geometry and facies, we interpret the amber-bearing shale of the El Soplao outcrop to have accumulated in a restricted tidal channel with low circulation and anoxic bottom-water, as suggested by widespread early pyritization.

3 Methods 3.1 Paleontological excavations Three paleontological excavations have been carried out in El Soplao amber outcrop. During the first one (October 2008) different extraction methods were applied to obtain amber pieces. Amber was obtained manually with small tools. In addition a large prospect hole, approximately 7 Ă&#x2014; 2.5 Ă&#x2014; 2 m in

size, was dug in one of the amber-rich areas of the outcrop using a bulldozer. Several tons of amber-bearing sediment from the large prospect hole were transported to a washing area located in the same outcrop, where a cement mixer and a sieve were used to obtain all ranges of amber sizes as described in Corral et al. (1999) and Alonso et al. (2000). This permitted to obtain a sampling without taphonomic biases introduced by the extraction methods towards the large and medium sizes. During the last excavation (July 2009) additional small prospect holes were dug close to the limits of the outcrop to find lateral extensions of the amber deposit. This task showed a high abundance of amber and revealed that the amber bed is at least 75 m long laterally, supporting the assertion of Najarro et al. (2009), which is that the El Soplao is the largest site of amber with arthropod inclusions that has ever been found in Spain so far.

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Fig. 2. Chrono- and lithostratigraphy of the El Soplao area. Modified from Hines (1985) and Najarro et al. (2010). Chronostratigraphy after Gradstein and Ogg (2009).

During the second paleontological excavation (March 2009) a new method was used to obtain abundant amber material increasing the collection of bioinclusions. The technique involves the use of high pressure water (Fig. 3.1 and 3.2) to extract entire large amber pieces that would be too fragile to resist conventional methods (Fig. 3.3). The water at high pressure disintegrated the sediment of the prospect hole and exposed the amber pieces, which were manually extracted to avoid fracturing. On the other hand, the small and medium pieces were retained using a large sieve where water and mud flowed (Fig. 3.2). The pieces from the mixture of mud, plant cuticles, coal and amber retained by the large sieve were separated in the washing area described above. This high pressure water method used for the first time to extract amber is already used in mining to fragment mineral seams, and to extract the useful parts from the waste rock. In conclusion, the new method was useful in obtaining numerous large amber pieces in a short period of time. 3.2 Biogeochemical analyses For the biogeochemical analysis amber pieces, fossil wood, and sediments rich in plant cuticles were collected from the El

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Soplao deposit during the first excavation on October 2008. Two types of amber pieces were collected: type A, characterized by a strong blue purple color under natural light (Fig. 3.3 and 3.4), purple-reddish under artificial light, and type B, less abundant, yellow under artificial light and with a bluish tinge under natural light. Plant cuticles were obtained from claystones by rinsing the plant-rich sediment in an ultrasonic bath of distilled water to remove all of the clay and silt sediment. The organic residue was air-dried. Plant fragments and leaves were distinguished and separated under a stereomicroscope. In addition, several resin and leaf samples from extant conifers of the families Cupressaceae and Araucariaceae (Cupressus arizonica, Agathis australis and Araucaria angustifolia) were collected from living trees at the Royal Botanic Garden of Madrid and in the kauri forests of New Zealand, in order to compare their compounds with those of the amber and fossil leaves and also to determine potential affinities of the amber with fossil taxa. For the analytical characterization, several representative pieces of amber of the types A and B of about 50 g each, with the highest transparency available and free of major inclusions, crusts and debris, were selected from the El Soplao deposit. Following standard techniques, each piece was crushed and extracted for 4 h with dichloromethane:methanol (2:1) using a Büchi model B-811 automatic extractor. One aliquot of extract was injected directly into the injection port of the gas chromatograph. The bulk extract was concentrated to a volume of 20 mL and fractionated by use of flash chromatography on silica gel. The elution was carried out using hexane, dichloromethane, dichloromethane:methanol (1:1), and methanol, and subsequently 25 fractions were collected. Each fraction was concentrated by evaporation of the solvent under N2 and analyzed by gas chromatography/mass spectrometry (GC/ MS). The fractions with similar compositions were combined. The polar fraction (eluted with methanol) and the fractions containing ferruginol were recombined, further separated using a glass column (20 cm) filled with chromatographic-grade silica gel, and eluted sequentially with n-hexane:dichloromethane (1:1), pure dichloromethane, dichloromethane:methanol (1:1), and methanol. Four fractions were collected, designated A through D. All fractions were dried and the alcohols and acids converted to trimethylsilyl derivatives by reaction with N,O-bis(trimethylsilyl)trifluoroacetamide (BSTFA) containing 1% trimethylchlorosilane (TMCS) at 65ºC for a period of 3 h. Finally, the derived fractions were diluted with n-hexane and injected into the port of the gas chromatograph. To study the molecular composition of fossil Frenelopsis and Arctopitys leaves (genus name Mirovia changed to Arctopitys; see Nosova and Wcisło-Luraniec, 2007), 5 g of leaves were extracted for 4 h with dichloromethane:methanol (2:1) using a Büchi model B-811 automated extractor. The bulk extract was filtered and analyzed directly by gas chromatography-mass spectrometry. After, extract was fractionated using silica gel chromatography in two fractions by elution with hexane:dichloromethane (3:1) and dichloromethane:methanol (4:1). The polar fraction was dried and derived using the method described above. For comparison, the chemotaxonomy of extant Cupressaceae and Araucariaceae was

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Fig. 3. Paleontological excavation in El Soplao amber outcrop (March 2009) with a new extractive method and some of the pieces obtained in situ. 1: High pressure water directed to the prospect hole; 2: large sieve to which water and mud flowed; 3: flattened amber piece virtually complete exposed in situ by high pressure water showing intensive fluorescent blue glow on its fracture (coin diameter 24 mm); 4: fragment of amber in situ showing intensive fluorescent blue glow (coin diameter 23 mm); 5â&#x20AC;&#x201C;6: part and counterpart of four-time branched shoots of Frenelopsis sp. collected during the last excavation carried out (July 2009).

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investigated using resin samples from Cupressus arizonica, Agathis australis, and Araucaria angustifolia. The resin was dissolved in dichloromethane:methanol (1:1) and fractionated by chromatography on silica gel in two fractions: the less-polar fraction being eluted with n-hexane:dichloromethane (1:4), and the second polar fraction being eluted with methanol. The polar fraction was dried under a nitrogen stream, yielding a white powder composed mainly of resin acids and highly polar compounds. The polar components were converted into trimethylsilyl derivatives by BSTFA and analyzed using GC-MS as described previously (Menor-Salván et al., 2010). 3.3 Palynological method Two samples, Sop-Peñosas (this from the amber outcrop) and Peñosas-Cóbreces, both from Las Peñosas Formation, were prepared for palynological studies in the laboratory of ALICONTROL (Madrid, Spain). The rock samples were treated following the standard palynological preparation technique (Batten, 1999), which consists of an acid attack with HCl, HF and HNO3 at high temperature. The residue was concentrated and sieved throughout sieves of different grid sizes (500, 250, 75, 50 and 12 μm). Then, the samples were mounted in glycerin jelly on glass slides for light microscopy. The samples were studied with an Olympus BX51 optical microscope. Both samples yielded representative and well-preserved assemblages: Sop-Peñosas yielded 488 miospores and Peñosas-Cóbreces 681 miospores. 3.4 Microscopic photography Scanning electron micrographs of the charcoal were taken using a HITACHI model S-2500 of the University of Valencia. Optical photography used both a digital camera attached to a microscope Olympus BX51 and a digital camera Leica DFC420 attached to a stereomicroscope Leica MS5.

4 Biogeochemistry of the Amber The overall aim of this part of the research was to identify the bioterpenoids preserved in the fossil resin and to determine their possible botanical source. Due to exceptional preservation, the amber-bearing deposit at El Soplao offers a unique opportunity to compare the molecular composition of the amber with that of the plant remains that appear in the same deposit. The analysis of the polar terpenoids of the amber from El Soplao indicates that most likely at least two resin producers contributed to the amber record (Fig. 4). The main parent resin (type A; Fig. 4.1) originally contained phenolic abietanes (dominated by ferruginol), totarol, dehydroabietane and pimarane/isopimarane acids. The dominant resin acids found were 13-dihydroagatholic and bisnordehydroabietic acids, as well as various other alteration products and minor quantities of callitrisic acid and hinokiol. The second parent resin (type B; Fig. 4.2) shares some general compounds characteristic of conifers with type A amber, but shows remarkable differences in specific biomarkers. It contained pimarane/isopimarane acids as the only identifiable biological precursors preserved and shows

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absence of phenolic terpenoids (ferruginol, totarol, hinokiol) and other specific biomarkers that are present in type A amber (e.g. dehydroabietane, callitrisic acid). The direct diagenetic products of the pimarane/abietane and labdane class terpenoids constitute the main geoterpenoids extractable from both types of El Soplao amber. The moderate degree of burial diagenesis of the studied material deduced from the low vitrinite reflectance values is consistent with the high level of preservation of the natural product diterpenoids and their direct diagenetic derivatives. The preliminary molecular data obtained from the comparative study of the megafossil plant leaves, extant plant samples and the two amber types lead to the following results: (a) There is absence of abietic and dehydroabietic acids in both types of amber samples. This allows us to reject a relationship between the amber and resin of Pinaceae species. Also, the absence of triterpenoids and labdatriene acids discards the contribution of angiosperms (Anderson et al., 1992; Yamamoto et al., 2006). (b) In type A amber, the phenolic diterpenoids present in the analyzed samples and the absence of phyllocladane/kaurane type terpenoids discard the contribution of the family Araucariaceae. The presence of phenolic terpenoids (ferruginol, totarol and hinokiol) points to a relation with the extant conifer families Cupressaceae, Taxodiaceae and Podocarpaceae. However, the presence of dehydroabietane (also present in representatives of Pinaceae and Cupressaceae; Otto et al., 2007) points to a relationship with the family Cupressaceae. The presence of callitrisic acid in type A amber reinforces a biochemical relation between the parent resin of amber and modern Cupressaceae, because in modern conifer resins the synthesis of callitrisic acid seems to be restricted to certain genera of Cupressaceae (Anderson, 2006). The analysis of the molecular composition of fossil leaves from the same outcrop shows the presence of key terpenoids, such as ferruginol, in both type A amber and the analyzed Frenelopsis fossil leaves, suggesting that this amber could be derived from the genus Frenelopsis (Cheirolepidiaceae). For the type A amber, a possible diagenetic route is suggested in Fig. 5 that connects the preserved biological precursors and the major geoterpenoids found in the sample (Otto and Simoneit, 2002; Stefanova et al., 2002; Hautevelle et al., 2006; Pereira et al., 2009). Morphological similarities between extinct Cheirolepidiaceae and extant Cupressaceae has been already described, but their phylogenetic relationship remains speculative, mainly due to the lack of molecular evidence (Seoane, 1998; Miller, 1999; Farjon, 2008); there are also paleobotanical data that support close affinities between Cheirolepidiaceae and Araucariaceae mainly based on female cone morphology. The chemotaxonomical affinity between type A amber and the analyzed leaves of Frenelopsis (Cheirolepidiaceae), and the affinity between type A amber and Cupressaceae as well, strongly suggests biochemical affinity between the extinct Frenelopsis and the modern representatives of the family Cupressaceae. (c) The overall terpenoid composition of B samples (Fig. 4.2) is represented by non-specific conifer biomarkers. Absence of

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Fig. 4. Gas chromatography-mass spectrometry traces (TIC) analysis and main biomarkers identified in the two types of amber found at El Soplao deposit: type A (a) and type B (b). After Menor-Salvรกn et al. (2010).

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Fig. 5. Diagenetic pathways proposed for the El Soplao amber. Dotted box: biological precursors. After Menor-Salván et al. (2010). phenolic terpenoids and 13-dihydroagatholic acid, together with a major presence of diagenetic products of pimarane type terpenoids, saturated and unsaturated norabietanes and alkyltetralines in B samples, points to a different biological origin. Paleobotanical source for this type of amber could not be identified on the basis of the biomarker composition found so far. Although these results are preliminary and might change when more data become available from analyses of more samples, it is estimated that they already provide a good indication of the potential plant source that originated the El Soplao amber. Future analyses could include fossil wood, leaf cuticle of Nehvizdya, more amber samples and conifer resin samples of other taxa.

5 Palynology This preliminary palynological study complements that based on plant cuticles from the amber outcrop (see Najarro et al., 2009), revealing some additional aspects of the regional paleobotany. The studied samples from Las Peñosas Formation

have yielded 32 spore and 29 pollen types (Fig. 6; see Table 1). In addition, undetermined bad-preserved dinoflagellate cysts and lining of foraminifera scarcely occurred. The Peñosas-Cóbreces sample presents more abundant taxa (52) than the Sop-Peñosas sample (44). It may be related to a lower palynological richness in the second sample. Both samples present high percentages of conifer pollen grains, especially those of Classopollis and Inaperturopollenites dubius. The spores of vascular cryptogams are more diverse than pollen grains although they occur in lower amounts. The most usual spores belong to both the trilete and taeniate genus Cicatricosisporites and the trilete and psilate genus Deltoidospora. Pollen grains of ancient angiosperms occur scarcely in both samples. They are mainly represented by monosulcate and reticulate pollen grains of the genus Clavatipollenites. The Peñosas-Cóbreces sample exhibits high amounts of pollen grains of Classopollis (~40%) and I. dubius (~30%). Other gymnospermous pollen grains such as Alisporites spp., Araucariacites australis, Monosulcites chaloneri and Spheripollenites sp. present lower but remarkable percentages.

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Fig. 6. Miospores from Peñosas-Cóbreces (1–5) and Sop-Peñosas (6) palynological samples. 1: Liliacidites dividuus (Pierce) Brenner; 2: Cicatricosisporites patapscoensis Brenner; 3: Densoisporites velatus Weyland & Krieger; 4: Taurocosporites segmentatus Stover; 5: Pennipollis peroreticulatus Brenner; 6: Gleicheniidites senonicus Ross.

Spores also present low amounts being Cicatricosisporites spp. the best represented (~2%). The occurrence of a single tricolpate pollen grain of the genus Tricolpites is noteworthy, because up to the Middle Albian this taxon is not abundantly represented. The Sop-Peñosas sample corresponds to the El Soplao amber outcrop. The palynological assemblage inferred in this sample is characterized by a conspicuous increase in I. dubius (~51%) and a marked decrease in Classopollis (~11%). The percentage of spores of the genus Deltoidospora also increases (~8% ). Cicatricosisporites scarcely occurs at this sample. Angiosperm

pollen grains are low represented in this sample compared to Peñosas-Cóbreces. From a biostratigraphical point of view, the occurrence of Appendicisporites robustus and Cicatricosisporites patapscoensis (Fig. 6.2) indicates a Late Aptian–Middle Albian age. However, the occurrence of Liliacidites dividuus (Fig. 6.1) and the low presence of Tricolpites sp. indicate an Early Albian age for the Peñosas-Cóbreces sample (Doyle and Robbins, 1977).

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Table 1 List of spores and pollen grains recorded from the Early Albian sediments of Las Peñosas Formation Spores Appendicisporites robustus Kemp 1970 Appendicisporites dentimarginatus Brenner 1963 Appendicisporites tricornitatus Weyland & Greifeld 1953 Appendicisporites spp. Baculatisporites sp. Biretisporites potoniaei Delcourt & Sprumont 1955 Ceratosporites sp. Cibotiumspora jurienensis (Balme 1957) Filatoff 1975 Cicatricosisporites apicanalis Phillips & Felix 1971 Cicatricosisporites apitereus Phillips & Felix 1971 Cicatricosisporites patapscoensis Brenner 1963 Cicatricosisporites recticicatricosus Döring 1965 Cicatricosisporites venustus Deák 1963 Cicatricosisporites spp. Cingutriletes sp. Contignisporites spp. Crybelosporites sp. Deltoidospora australis (Couper 1953) Srivastava 1975 Deltoidospora minor (Couper 1953) Pocock 1970 Deltoidospora sp. Densoisporites velatus Weyland & Krieger 1953 Dictyophyllidites harrisii Couper 1958 Echinatisporis sp. Gleicheniidites senonicus Ross 1949 Laevigatosporites sp. Leptolepidites sp. Patellasporites tavaredensis Groot & Groot 1962 Retitriletes sp. Stereisporites sp. Taurucosporites segmentatus Stover 1962 Trachysporites sp. Triporoletes reticulatus (Pocock 1962) Playford 1971 Todisporites major Couper 1958 Pollen grains (gymnosperms) Alisporites bilateralis Rouse 1959 Alisporites spp. Araucariacites australis Cookson 1947 Callialasporites dampieri Dev 1961 Cedripites sp. Classopollis classoides Pflug 1953 emend. Pocock & Jansonius 1961 Classopollis spp. Cycadopites spp. Eucommiidites minor Groot & Penny 1960 Exesipollenites tumulus Balme 1957 Ginkgocycadophytus nitididus (Balme 1957) de Jersey 1962 Inaperturopollenites dubius (Potonié & Venitz 1932) Thompson & Pflug 1953 Inaperturopollenites spp. Monosulcites chaloneri Brenner 1963 Monosulcites sp. Pinuspollenites sp. Podocarpidites sp. Spheripollenites sp. Vitreisporites pallidus (Reissinger 1950) Nilsson 1958 Undetermined bisaccate pollen grains Pollen grains (angiosperms) Afropollis sp. Clavatipollenites hughesii Couper 1958 Clavatipollenites minutus Brenner 1963 Clavatipollenites sp. (trichomosulcate) Clavatipollenites spp. Liliacidites dividus (Pierce 1961) Brenner 1963 Pennipollis peroreticulatus Brenner 1963 Tricolpites sp. Undetermined angiospermous pollen grains TOTAL miospores

Peñosas-Cóbreces

Sop-Peñosas

1 0 1 2 0 3 1 0 1 1 1 1 2 13 1 1 0 4 8 4 1 2 1 0 0 0 1 0 0 1 1 1 1

0.147 0 0.147 0.290 0 0.440 0.147 0 0.147 0.147 0.147 0.147 0.290 1.908 0.147 0.147 0 0.587 1.174 0.587 0.147 0.290 0.147 0 0 0 0.147 0 0 0.147 0.147 0.147 0.147

0 2 0 1 1 3 0 1 0 0 0 0 1 6 1 0 1 13 27 8 1 1 0 4 7 1 0 1 1 0 0 1 0

0 0.410 0 0.205 0.205 0.615 0 0.205 0 0 0 0 0.205 1.230 0.205 0 0.205 2.664 5.533 1.640 0.205 0.205 0 0.819 1.434 0.205 0 0.205 0.205 0 0 0.205 0

5 20 12 1 1 213 58 8 4 6 1 205 21 22 4 2 0 12 2 1

0.734 2.937 1.762 0.147 0.147 31.277 8.517 1.174 0.587 0.881 0.147 30.102 3.084 3.230 0.587 0.290 0 1.762 0.300 0.147

8 12 25 0 0 46 9 2 2 3 3 249 9 8 3 2 2 5 1 6

1.640 2.460 5.123 0 0 9.426 1.844 0.410 0.410 0.615 0.615 51.024 1.844 1.640 0.615 0.410 0.410 1.024 0.205 1.230

2 2 5 4 5 3 1 1 6 681

0.300 0.300 0.734 0.587 0.734 0.440 0.147 0.147 0.881 100

1 3 1 0 4 0 1 0 1 488

0.205 0.615 0.205 0 0.819 0 0.205 0 0.205 100

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The area could have been covered by mixed conifer forests of Cupressaceae and Cheirolepidiaceae that grew near the sea. Their understory was integrated by pteridophytes, cycads and/or Bennettitales. Ponds and swampy areas were mainly occupied by vascular cryptogams and early angiosperms, which could have aquatic habits. The predominance of Classopollis and the lower amount of spores observed in the Peñosas-Cóbreces sample could be related to a drier period. The more humid conditions of Sop-Peñosas are indicated by the higher percentages of I. dubius and Deltoidospora spp. as well as those of Laevigatosporites sp., Alisporites spp. and Araucaricites australis. Considering their composition, the assemblages are similar to those from the Upper Aptian–Lower Albian sediments of the Oliete sub-basin (Iberian Ranges) (Peyrot et al., 2007a, 2007b).

6 Plant Cuticles Najarro et al. (2009) reported abundant plant cuticles in the amber-bearing beds of El Soplao outcrop, sometimes as levels up to 10 cm thick. This plant assemblage comprises female cones of the genus Alvinia, leafy axes of Brachyphyllum-type, Nehvizdya sp. (and its reproductive organs classified into the genus Nehvizdyella), Pseudotorellia sp., and mainly Frenelopsis and Arctopitys (cited as Mirovia in that paper, but see Nosova & Wcisło-Luraniec, 2007). Up to four-time branched shoots of the cheirolepidiaceous conifer Frenelopsis have been collected during the last excavation in July 2009 (Fig. 3.5 and 3.6), due to the large area exposed in the prospect hole excavated. The shoots show lateral branches with a single branch per node, a branching pattern similar to that described and discussed by Daviero et al. (2001). These new records are of taphonomic and paleoenvironmental relevance. From the taphonomic point of view, the stiff, articulated leafy internodes of Frenelopsis were probably very brittle and fragmented when transport occurred (Gomez et al., 2001, 2002; Riera et al., 2010). At least for some of the plants of the insect-bearing amber assemblage of El Soplao, such an organization suggests a parautochthonous deposition. The representatives of Frenelopsis occupied habitats from freshwater wetlands to saline, coastal or estuarine marshes (e.g. Gomez et al., 2001, 2002; Mendes et al., 2010). Such a wide range of habitats suggests possible mangrove-like ecology for the Frenelopsis of El Soplao. This paleoecological inference is also supported by the sedimentological context and the occurrence of marine or brackish-water invertebrates in the sediment and on amber surfaces.

7 Charcoal and Charcoalified Plant Fibers The first record of charcoalified (=fusainized) plant remains indicating paleofires is from the Silurian from Ludford Lane in the Welsh Borders in England (Glasspool et al., 2004). Scott (2000) provided a general view of the Pre-Quaternary history of fire based on charcoalified plant remains, which are especially important in sediments from the latest Jurassic and Early Cretaceous. Charcoalified plant remains are moderately common

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in a wide variety of facies, and some notable concentrations of considerable paleobotanical and paleoecological value can occur locally (Nichols et al., 2000). In Spanish outcrops, amber and charcoal are very scarce except for a few levels in which both appear abundantly. This is the case of all the main outcrops for which this aspect has been explored: Peñacerrada (Suárez-Ruiz, 2003 and López del Valle, per. obs., 2008), La Hoya, San Just (Peñalver et al., 2007, 2008) and El Soplao. In these outcrops there are abundant cuboidal pieces of charcoal (=fusinite) of a few centimeters long (Fig. 7.2), and they are easily recognized with the naked eye even in the field by its silky luster, brittleness and friability. In addition, charcoal pieces are fibrous, black and opaque. Features of the charcoal associated with amber in El Soplao outcrop using SEM are the undeformed structure and fabric of the tissue with open cell lumina, pits and homogenization of the cell walls (Fig. 7.3–7.5); see Sander and Gee (1990) and Scott (2000, 2010) for general description of charcoal. All of these features revealed by SEM examination are particularly characteristic of a pyrolysis origin. The most important feature is that different layers of the cell wall and adjacent cell walls cannot be distinguished from one to another due to the homogenization during burning. This homogenization of the wall takes place above 300ºC and the resulting charcoal is highly resistant to microbiological or chemical degradation during sedimentation and diagenesis (Cope and Chaloner, 1980). In San Just and El Soplao outcrops charcoal pieces with the cell lumina diagenetically filled with framboidal pyrite are found frequently (Fig. 7.6). The levels of concentration of amber and charcoal in Spanish outcrops, including El Soplao, reveal paleofires in the resinous forests. Grimaldi et al. (2000) reported wood, insect remains and flowers, all charcoalified, and fire-damaged amber from the Upper Cretaceous (Turonian) of New Jersey, and Jarzembowski et al. (2008) reported a single piece of amber with similar features from the Lower Cretaceous of the Isle of Wight. Brasier et al. (2009) reported abundant examples of amber associated with charcoalified wood from the Early Cretaceous (140 Ma) amber deposit in Hastings (Sussex). In addition to the frequent presence of charcoal associated with the amber in the same beds, a few amber pieces from El Soplao contain charcoalified plant fibers that appear dispersed inside the amber (Fig. 7.1). It is the first time in the fossil record that charcoalified fibers as bioinclusions have been reported. These fibers are small (around 0.7 mm long) and can be recognized as charcoalified fibers due to both their opaque black color and silky luster. This appearance is not due to the fossildiagenetic process because they contrast to other plant fibers in the same amber pieces that have the common translucent clear brown color without silky luster. These small charcoalified plant fibers became included in resin during paleofires moved by convective currents or, after paleofires, transported by the wind to the exposed resin from the soil of burned areas. Scott et al. (2000) reported how abundant finer charcoal material was transported by wind from a charred area of the Frensham Common Country Park (England) several days

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Fig. 7. Fossil carbonized woody tissues (charcoals) from El Soplao outcrop. 1: Three charcoalified plant fibers from the same small amber piece containing numerous of these remains (at the same scale); 2: cuboidal piece of charcoal; 3–5: scanning electron micrographs of charcoals show both pits and cell-wall homogenization in epidermis and peripheral tissues (micrograph 4 is a detail of 3); 6: framboidal pyrite partially filling the cell lumina (micrographs 3–6 are from microsamples of the piece figured in picture 2).

after a wildfire. The kauri forests (Araucariaceae) of New Zealand most likely are the best extant correlate to the Cretaceous resiniferous forests, except, perhaps, for the fact that the kauri forests do not have a fire ecology (Daniel J. Bickel, pers. comm., 2010). As it

has been observed by us in the kauri forests, very large amounts of organic material occur due to the floor accumulation of litter, including resin pieces; Silvester and Orchard (1999) indicated that litter around a large kauri tree may reach 2 m or more in depth, with a mean residence time of 9–78 years due to a slow

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decomposition rate. Although resin burns easily, forest fires only affect the litter superficially, except for ground fires (Scott, 2000). Thus the main part of the accumulated resin usually remains intact. Scott et al. (2000) reported a surface fire in the Frensham Common Country Park that charred only a few millimeters of the organic litter, so the pine crown was practically not affected by the fire. As indicated by Grimaldi et al. (2000) and Martínez-Delclòs et al. (2004), forest fires traumatically induced copious production of resin and they might have been an important factor in the genesis of amber deposits. From this respect, Brasier et al. (2009) reported charcoalified conifer wood from Hastings deposit with cell lumina filled with resin after the paleofire. However, we consider that the joined occurrences of large accumulations of amber and charcoal in Spanish outcrops, and abundant plant cuticles as well, were mainly the consequence of an intensive erosion of the litter in burnt forests, possibly including mass wasting; at this respect, Scott & Stea (2002) reported evidences of post-fire soil erosion in Cretaceous charcoal horizons of Nova Scotia. As summarized by Nichols et al. (2000) from different sources, the removal of vegetation by modern fires can increase erosion rates by weathering and during rainstorms by up to 30 times compared with the sink of pre-fire levels, and single wildfire events may be recognizable as responsible for individual depositional units, for example on alluvial fans. Nichols et al. (2000) conducted actualistic experiments and concluded that charcoal is an unusual sedimentary material because most fresh material floats, but after prolonged immersion becomes waterlogged and sinks, mainly the small pieces. After the study of an area of heathland in SW England burned by an uncontrolled fire, Blackford (2000) concluded that large particles are not transported long distances and thus are indicative of local fires. The great abundances of centimetersized charcoal in amber-bearing beds of San Just and El Soplao outcrops suggest that paleofires in these cases were local or occurred close to the area of deposition. In the scenario we propose, intensive erosion by rivers and storm floods of a burned resinous forest area located close to the deltaic environments of deposition was promoted by the loss of vegetation, and resin plus charcoalified wood were transported together through water to coastal environments. Another complementary factor that could contribute to the great amber accumulation in El Soplao outcrop was proposed by Najarro et al. (2009) based on sedimentological data: floods during rainstorms eroded and removed the resin and plant remains from the soils of coast-fringing forests. Pieces of resin and wood mixed with mud and sand then were transported by density flows to the coastal and interdistributary bays. Surely, more than one of these three factors might have occurred. We can assume that wildfires had a great impact on environments and most likely they occurred mainly during the warm, drought season; lightning strike is the most important reported cause of naturally ignited fires today (Cope and Chaloner, 1980; Scott, 2010) and it was probably the same during the Cretaceous. Moreover, levels of atmospheric oxygen during the Cretaceous were among the highest during the

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Phanerozoic, close to 30% (vs. 20% today), leading to a much higher prevalence of fires (Robinson, 1989). Secondary evidence of the environmental impacts of wildfires can be obtained from plant and insect fossil records. The tree fern Weichselia reticulata has been recorded in some adjacent beds to the Spanish amber outcrops and this taxon can be related to environments disturbed by both floods and fires according to Coiffard et al. (2007) and Scott et al. (2000). More recently Ortega-Blanco et al. (2008) described a new species of anaxyelid woodwasp, Eosyntexis parva, from Peñacerrada I amber closely related to the extant anaxyelid taxa that lay eggs in burnt conifer trees shortly after the wildfires.

8 Marine or Brackish-water Invertebrates on Amber Surfaces El Soplao amber deposit in some levels contains internal marcasite moulds of marine or brackish-water mollusks, and oyster shells that preserve the original calcium carbonate biomineralization. In addition, for the first time in the fossil record we report findings of large amber pieces that show their surfaces colonized by both serpulid worm tubes (Annelida: Polychaeta) (Fig. 8.1 and 8.2) and encrusting bryozoan colonies (Bryozoa: Cheilostomata) (Fig. 8.3); barnacles encrusting Baltic amber surfaces have been reported, but they are recent barnacles (see Grimaldi, 1996). Amber pieces with their entire surface colonized, or only with one side colonized, have been found. Serpulid worm tubes exhibit four modes of fossilization that sometimes occur in the same amber piece, but in different parts: (1) as limonitic impressions of the tubes (Fig. 8.1); (2) as internal marcasite moulds (Fig. 8.2); (3) as amber tubes (internal fillings with resin during first fossildiagenetic phases); and (4) as original carbonate remains (Fig. 8.3). Serpulid tube remains are several millimeters long and internal tube fillings are approximately 0.35 mm diameter. Bryozoan colonies left marks of the zooid exoskeletons on the amber surface and when original calcium carbonate remains are preserved (Fig. 8.3) they easily fade away during the washing of the pieces. Bryozoan colonies are multiserial and belong to the Order Cheilostomata; that group produces mineralized exoskeletons and form single-layered sheets that encrust over surfaces. Box-shaped, rectangular zooids are 0.34 × 0.29 mm in size. Completeness of these large pieces suggests that they were constituted by polymerized resin when serpulids and bryozoans developed, not fossil resin from eroded strata (reworking process). Resin is comparatively less fragile than amber or copal and, in consequence, more commonly keeps its integrity during marine transport. In any case, this deposit contained a mixture of resin pieces from the litter, and fresh resin pieces directly transported from the trees, together with reworked resin pieces that remained a certain time in saline water. The presence of encrusting bryozoan colonies on the amber surfaces is not surprising because this type of organism is better adapted to shallow, high energy environments. In conclusion, this exceptional record indicates a littoral to coastal marsh

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Fig. 8. Marine or brackish-water invertebrates on the surfaces of El Soplao amber and insects as bioinclusions. 1: Detail of the serpulids preserved as limonitic impressions on the entire kidney-shaped amber piece 12 Ă&#x2014; 10 Ă&#x2014; 5 cm in size; 2: two serpulid worm tubes preserved as marcasite internal mould of the same amber piece; 3: bryozoan colony and several serpulids preserved as carbonate fossils on the surface of a large amber piece; 4: virtually complete specimen of the family Berothidae (Neuroptera); 5: specimen of Cucujoidea (Coleoptera). Photos 4 and 5 were made with integrated consecutive pictures taken at successive focal planes.

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environment of deposition for El Soplao outcrop.

9 New Data on Insect Bioinclusions Najarro et al. (2009) reported a high abundance of arthropod specimens embedded in the Early Cretaceous amber of El Soplao outcrop, and explained it as a consequence of the unusual concentration of amber pieces indicative of resin flows (drops, crusts and runnels) in this deposit. A significant percentage of amber pieces of this type contains numerous bioinclusions because the original resin was exposed to the atmosphere, was less viscous (it was more easily penetrated by insects), and successive flows encapsulated the trapped insects (MartínezDelclòs et al., 2004). Recent paleontological excavations have provided new bioinclusions of significant paleoecological and taxonomical value. To date, El Soplao amber has provided more than 200 bioinclusions, including fungi, plants, and diverse arthropods. Among them, only systematic studies on insects have been already started. The insect inclusions found belong to 11 recognized orders: Blattaria, Isoptera, Psocoptera, Thysanoptera, Raphidioptera, Neuroptera, Hemiptera, Coleoptera, Trichoptera, Hymenoptera, and Diptera; last two groups are the most abundant. Neuropterid fauna from El Soplao amber is currently being studied. It is composed so far of Neuroptera and Raphidioptera, including the new genus and species of mesoraphidiid described by Pérez-de la Fuente et al. (2010). Neuropterans are represented by two specimens of dustywings (Coniopterygidae) and two beaded lacewings (Berothidae), one of them very complete and interpreted as a new morphotype (Fig. 8.4). Surely, the most outstanding finding is a tentative green lacewing larva (Chrysopidae), which, if confirmed, would be the oldest representative of the family in amber and one of the very few fossil chrysopid larvae ever reported. Its morphology, to be described and discussed elsewhere, is similar to those of the extant trash-carrying chrysopid larvae. Coleopterans from El Soplao amber are represented only by members of the suborder Polyphaga, and all of them are up to now interpreted as herbivorous or saproxylic forms, most probably living under the bark of the trees. As in other Lower Cretaceous outcrops of surrounding areas in Spain and France, the most abundant forms of beetles from El Soplao belong to superfamilies Elateroidea and Cucujoidea (Fig. 8.5), but there are also present forms of Curculionoidea (?Attelabidae, ? Nemonychidae) and possibly a member of Dascilloidea. The presence of two specimens of Curculionoidea in such small collection of beetles (approximately 10 specimens) is of paleoecological relevance, because this group of beetles in extant ecosystems is closely related to hard parts of trees, in which larval stages developed, and the action damaging trees has also been pointed to as one of the possible causes for the increase in resin-production in trees during the Early Cretaceous and its subsequent record as fossil resins. Although the hymenopteran assemblage of El Soplao is scarce, it is providing interesting data. It is composed of

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parasitoid wasps of several superfamilies as Ceraphronoidea (parasitoids of numerous taxa, hyperparasitoids sometimes), Platygastroidea (insect and spider egg parasitoids), Evanioidea (cockroach egg consumers during larval stage), Ichneumonoidea (commonly hyperparasitoids) and Mymarommatoidea (tiny egg parasitoids). Although ceraphronoids are currently under study, a first approximation allows us to recognize members of a recently described new family of Ceraphronoidea (Ortega-Blanco et al., 2010b). In Spanish amber, Ceraphronoids were only detected in the Peñacerrada amber until date; venation and some antennae aspects of the El Soplao morphotype indicate that it belongs to the same genus that is present in Peñacerrada amber. Platygastroids are the most common hymenopteran insects in El Soplao amber, as in all other Spanish ambers, and the study of this group is in progress. Apparently platygastroids belong to a new species with large, 14-articled antennae. A new parasitoid wasp species of the genus Archaeromma (Mymarommatoidea) has been recorded in El Soplao, as it was in Peñacerrada I. Recording the same insect species supports the hypothesis that the age of these ambers was equivalent or at least very close, and suggests that both outcrops had similar paleoecological characteristics. Dipterans from El Soplao amber are represented by Brachycera and Nematocera forms. One specimen of Litoleptis sp. (Spaniidae) has been identified within the Brachycera, similar to one specimen described from San Just amber (Arillo et al., 2009). Other dipteran families present include Hybotidae (Brachycera), as well as Cecidomyiidae and Psychodoidea. The current paleodiversity of biting midges (Ceratopogonidae) has been evaluated, leading to the discovery of five species, including three new species (Najarro et al., 2009). The new species belong to the genera Archiaustroconops, Szadziewski, 1996, Atriculicoides Remm, 1976, and Lebanoculicoides Szadziewski, 1996. All these ancient biting midges would have shown a hematophagous diet which is considered plesiotypic within the family (Borkent, 1995).

10 Conclusions We present herein data that suggest that amber pieces from El Soplao have at least two botanical sources. Part of these data strongly supports a source related to Cheirolepidiaceae and also suggests a molecular relationship of Frenelopsis with extant Cupressaceae. This relationship is based on the chemotaxonomic comparison of biological diterpenes found in amber and plant megafossils and extant Cupressaceae. The taphonomic significance of wildfires in the origin of the Cretaceous amber accumulations has received little attention and has been clearly underestimated. The potential presence of charcoal, fire-damaged amber pieces and charcoalified plant fibers embedded inside the amber must be considered when studying the origin of the Cretaceous amber deposits. In fact, the occurrence of large accumulations of amber and charcoal in Spanish outcrops were consequence of intensive erosion of the litter in burnt forests, possibly including mass wasting. This scenario has been well documented in the case of the El Soplao

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deposit, but similar studies on other Spanish amber deposits should be conducted. The new data on the biological inclusions from El Soplao amber show a highly diversified entomofauna, mainly represented by dipterans and hymenopterans, with the presence of very specialized groups such as chrysopids or weevils, which may help to better understand the early evolution of some insect groups. The data previously published and new data herein provided suggest a taphonomic history for the El Soplao amber that may have run approximately as follows: resin was exuded by conifers (perhaps by two different conifers, including Cheirolepidiaceae) closely to the deltaic environments of deposition. Conifer forests were well-present in the region, with understory integrated by pteridophytes, cycads and/or Bennettitales; ponds and swampy areas were mainly occupied by vascular cryptogams and early angiosperms, which could have aquatic habits. A diversified insect fauna, mainly represented by hymenopterans and dipterans, developed around the conifer trees and was abundantly embedded in resin. Coleopterans, which probably lived under the bark or in close contact with the wood of the conifer trees, were abundantly embedded as well. Wildfires in the resinous forests promoted both the resin production and an intensive erosion of the partially burned litter. Resin, fresh leaves, wood and charcoal were shortly transported together through water and accumulated alongside marine or brackish-water mollusks in a restricted tidal channel with low circulation and anoxic bottom-water. The conifer Frenelopsis, a potential source of resin as suggested by the biogeochemical analyses, grew close to the deposition area, where articulated branched shoots accumulated (parautochthonous deposition). Resin pieces differed in their biostratinomic histories because some of them remained a certain time in shallow, high energy saline waters where serpulids and bryozoans grew on the resin surfaces, resulting in a mixed assemblage. Low circulation and anoxic bottom-water promoted an early pyritization of the marine mollusks, charcoal and amber surfaces, including the fixed serpulid tubes. The assemblage was lastly buried by siltstones and sandstones. The maximum temperatures suffered by the amber deposit during diagenesis were in the range of 60–70ºC. The inferred low maturity levels were maybe responsible for the good preservation of amber molecular composition and its biological inclusions.

Acknowledgements This work is part of the Ph.D. Thesis of three of the authors (M.N. -geology-, R.P.F. and J.O.B. -paleobiology-), which are supported by a scholarship from the Instituto Geológico y Minero de España (IGME), an APIF grant of the University of Barcelona, and a FPI grant from the Spanish Ministry of Science and Technology, respectively. This study is a contribution of the IGME Project 491-CANOA 35015 “Investigación científica y técnica de la Cueva de El Soplao y su entorno geológico”, the projects CGL2008-/01237BTE from the MICINN, CGL200800550/BTE: “Amber of the Cretaceous of Spain: A multidisciplinary study”, and the ANR Project AMBRACE

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BLAN07-1-184190. The contribution is framed in a collaborative agreement among the Cantabrian Government (Regional Cultural, Tourism and Sports Ministry), IGME and SIEC S.A. We thank Gonzalo Nieto and the staff of the Royal Botanic Garden of Madrid for permission and assistance in the sampling of extant conifers. We also express our thanks to Francisco Javier López Marcano (Regional Minister of the Cantabrian Government), José Pedro Calvo Sorando (IGME) and Fermin Unzué (manager of the El Soplao Cave) for their efforts and promotion of the study of the outcrop. We are grateful to Dr Daniel J. Bickel (Australian Museum, Sydney), Dr Vladimir Blagoderov (Natural History Museum, London) and an anonymous reviewer for careful reviews of the manuscript. Thanks are due to all people that participated in the paleontological excavations. Manuscript received Jan. 4, 2010 accepted April 7, 2010 edited by Fei Hongcai

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Author's personal copy Sedimentary Geology 263–264 (2012) 174–182

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Tracing organic compounds in aerobically altered methane-derived carbonate pipes (Gulf of Cadiz, SW Iberia) Raúl Merinero a,⁎, Marta Ruiz-Bermejo b, César Menor-Salván b, Rosario Lunar a, Jesús Martínez-Frías b a b

Departamento de Cristalografía y Mineralogía, Facultad de Ciencias Geológicas, Universidad Complutense de Madrid, Avda Complutense s/n, 28040, Madrid, Spain Centro de Astrobiologia, CSIC/INTA, associated with the NASA Astrobiology Institute, Ctra de Ajalvir km. 4, 28850, Madrid, Spain

a r t i c l e

i n f o

Article history: Received 2 December 2010 Received in revised form 12 September 2011 Accepted 15 September 2011 Available online 22 September 2011 Keywords: Gulf of Cádiz Aerobic degradation Carbonate pipes Lipids

a b s t r a c t The primary geochemical process at methane seeps is anaerobic oxidation of methane (AOM), performed by methanotrophic archaea and sulfate-reducing bacteria (SRB). The molecular fingerprints (biomarkers) of these chemosynthetic microorganisms can be preserved in carbonates formed through AOM. However, thermal maturity and aerobic degradation can change the original preserved compounds, making it difficult to establish the relation between AOM and carbonate precipitation. Here we report a study of amino acid and lipid abundances in carbonate matrices of aerobically altered pipes recovered from the seafloor of the Gulf of Cadiz (SW Iberian Peninsula). This area is characterized by a complex tectonic regime that supports numerous cold seeps. Studies so far have not determined whether the precipitation of carbonate pipes in the Gulf of Cadiz is a purely chemical process or whether microbial communities are involved. Samples from this site show signs of exposure to oxygenated waters and of aerobic alteration, such as oxidation of authigenic iron sulfides. In addition, the degradation index, calculated from the relative abundance of preserved amino acids, indicates aerobic degradation of organic matter. Although crocetane was the only lipid identified from methanotrophic archaea, the organic compounds detected (n-alkanes, regular isoprenoids and alcohols) are compatible with an origin from AOM coupled with bacterial sulfate reduction (BSR) and subsequent aerobic degradation. We establish a relation among AOM, BSR and pipe formation in the Gulf of Cadiz through three types of analysis: (1) stable carbon and oxygen isotopic composition of carbonate minerals; (2) carbonate microfabrics; and (3) mineralogical composition. Our results suggest that carbonate pipes may form through a process similar to the precipitation of vast amounts of carbonate pavements often found at cold seeps. Our approach suggests that some organic compound patterns, in combination with additional evidence of AOM and BSR, may help indicate the source of altered methane-derived carbonates commonly occurring in ancient and modern deposits. © 2011 Elsevier B.V. All rights reserved.

1. Introduction At submarine methane seeps, groups of archaea and bacteria carry out anaerobic oxidation of methane (AOM) and sulfate-dependent anaerobic oxidation of methane (AOM) (e.g., Hinrichs et al., 1999; Boetius et al., 2000; Orphan et al., 2001; Reitner et al., 2005). AOM increases alkalinity, inducing the precipitation of authigenic carbonates (Hovland et al., 1987; Ritger et al., 1987; Paull et al., 1992). Molecular ﬁngerprints (biomarkers) of these chemosynthetic microorganisms are well preserved in carbonate matrices, the characterization of which is a useful tool to establish the relation between carbonate precipitation and methane seepage (Peckmann and Thiel, 2004). However, post-depositional processes such as biodegradation and thermal maturation alter the biomarker inventory of the carbonates, obscuring the information they provide (Goedert et al., 2003; Birgel et al.,

⁎ Corresponding author. Tel.: + 34 1 3944959; fax: + 34 1 3944872. E-mail address: rmeriner@geo.ucm.es (R. Merinero). 0037-0738/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.sedgeo.2011.09.011

2006) and making it difficult to establish the role of methane and microbial activity in carbonate formation. Methane-enriched fluid venting is a widespread process on the continental slope of the Gulf of Cadiz, as reflected in the abundance of methane-related seafloor structures in this area. These structures include pockmarks (Baraza and Ercilla, 1996; Casas et al., 2003); mud volcanoes, some of which contain gas hydrates (Pinheiro et al., 2003; van Rensbergen et al., 2005; Niemann et al., 2006); and carbonate-mud mounds bearing cylindrical carbonate deposits or chimneys, crusts and slabs (León et al., 2006). Biomarkers of methanotrophic archaea have been isolated and identified on carbonate crusts and pavements collected on mud-volcanoes of the Gulf of Cadiz (Niemann et al., 2006; Stadnitskaia et al., 2008). Previous studies have used the term “chimneys” to refer to carbonate deposits with cylindrical or conical shapes (Díaz-del-Río et al., 2003). This term implies a vertical channel that projects into the water column during its formation. However, it is more likely that these carbonate deposits formed in the sediment and were later shaped by erosion, making the term “pipe” more appropriate.

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Early studies suggested that the most probable origin of carbonate pipes is AOM coupled with sulfate reduction, but explicit evidence is still lacking (Díaz-del-Río et al., 2003; León et al., 2006, 2007). Moreover, the possibility of purely chemical precipitation of the carbonate pipes, without a microbial contribution, cannot be excluded due to the complex tectonic regime governing the Gulf of Cadiz (Maldonado et al., 1999) and the evidence of hydrothermal imprints in mud-volcano fluids in the Gulf of Cadiz (e.g. Hensen et al., 2007). We consider these arguments necessary in order to exclude processes other than methane oxidation as possible origins of carbonate pipes. This would be an important contribution of our study, since previous works on tubular carbonate deposits (chimneys) in the Gulf of Cadiz have not clarified this question. Previous studies of the carbonate pipes in the Gulf of Cadiz showed that they were lying over the seafloor, exposed to oxygenated waters (Díaz-del-Río et al., 2003). More recently, Merinero et al. (2008) reported the oxidation of iron sulfides and pseudomorphic transformation into iron oxyhydroxides. Thus, aerobic degradation of organic matter originally preserved inside the pipes was expected, which would make it difficult to establish the relation between carbonate pipe formation and AOM by identifying biomarkers from methanotrophic microbes. Here we analyze a large collection of cylindrical carbonate pipes from the Gulf of Cadiz using a combination of geochemistry (analysis of organic compounds and 13C and 18O isotopic composition of carbonate minerals), mineralogy and petrology. The objectives of this research were to (1) determine the degree of aerobic alteration of the pipes; (2) verify the expected relation between AOM and the origin of the pipes; and (3) characterize the organic compounds preserved within the pipes and establish their possible relation with aerobic degradation of biomarkers from chemosynthetic microorganisms.

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Mediterranean Sea. The segment, referred to as the Gibraltar Arc, lies between the Gloria Fault and the western end of the Betic-Rifean orogenic belt (Fig. 1). This area shows a complex geodynamic history with various extensional stages, strike-slip and compression associated with the closure of the Tethys, the opening of the North Atlantic, and the African-Eurasian convergence since the Cenozoic (Maldonado et al., 1999). Water exchange between the Mediterranean and Atlantic in the region of the Strait of Gibraltar has been a crucial factor in the development of various mesoscale physical structures characteristic of the Atlantic Ocean seafloor (Criado-Aldeanueva et al., 2006; GarcíaLafuente et al., 2006; Mulder et al., 2006). Hydrocarbon-rich fluids are trapped in a thick sedimentary formation called the Olistostrome or Allochthonous Unit of the Gulf of Cadiz (Medialdea et al., 2004). This formation emplaced during the Middle Miocene and consists of Triassic evaporites and red beds and blocks of Cretaceous to Paleogene limestones (Maldonado et al., 1999). Migration and subsequent seepage of the fluids have occurred because of diapirism and a compressional regime, which affect the Olistostrome in response to the Africa-Eurasia plate convergence (Maldonado et al., 1999; Medialdea et al., 2009). As a consequence of these processes, many types of seafloor structures formed (León et al., 2006): pockmarks (Baraza and Ercilla, 1996; Casas et al., 2003); mud-volcanoes (Pinheiro et al., 2003; van Rensbergen et al., 2005; Niemann et al., 2006); and carbonate-mud mounds bearing cylindrical carbonate deposits, crusts and slabs (Díaz-del-Río et al., 2003; Merinero et al., 2008). The close relationship observed between tectonic structures and these hydrocarbon-derived features suggests that fluid venting is triggered by the formation of pressurized compartments beneath thrust structures, which provide conduits for hydrocarbonenriched fluids (Maestro et al., 2003; León et al., 2007). 2.2. Material recovery and data acquisition

2. Area descriptions, materials and methods 2.1. Geological settings The Gulf of Cadiz is located at the trending segment of the Eurasian-African plate boundary that extends from the Azores to the

Samples were taken from the Iberian continental margin of the Gulf of Cadiz during the oceanographic cruises Anastasya 2000 and 2001 aboard the research vessel Cornide de Saavedra. The study area was extensively surveyed with swath bathymetry, multi-channel and ultra-high-resolution seismic reﬂection, gravimetry, magnetism,

Fig. 1. Bathymetric map of the Gulf of Cadiz (modiﬁed from Merinero et al., 2009) showing the locations of the four sites where samples were collected (IB = Iberico mound, CN = Cornide mound, AR = Arcos mud-volcano, CO = Coruña mud-volcano). Locations of the mud-volcanoes and mud-mounds are from León et al. (2006) and González et al. (2009).

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underwater cameras, dredging, and gravity coring. A large number of mud volcanoes, mud ridges, crater-like pockmarks, and large sediment slides were mapped. The carbonate pipes were discovered during the Anastasya 2000 cruise while dredging an 870 m-deep, 120 m-tall carbonate mound known as Iberico (Díaz-del-Río et al., 2003). Most pipes were collected along or in close proximity to the main channels of the Mediterranean Outﬂow Water using rectangular benthic-type dredges. More than 200 carbonate pipes were recovered at depths ranging from 850 to 1100 m during the cruises in 2000 and 2001. Here we studied 13 carbonate pipes from four different sites (Table 1). 2.3. Mineralogy, geochemistry and stable C and O isotope composition We examined each specimen as a hand sample and in thin section (transversal and longitudinal views). From these slabs, we prepared thin sections for standard petrographic analysis. We made SEM and microprobe analyses on carbon-coated thin sections using a Philips XL20 scanning electron microscope with accelerating voltages of 20–30 kV. We determined the carbonate mineralogy by X-ray diffraction (XRD) using a Philips PX-1710 diffractometer operating at 40 kV, equipped with a graphite monochromator and an automatic divergent slip. We also analyzed the carbonate chemical composition by electron microprobe (JEOL Superprobe JXA-8900M). For isotopic measurements, samples were ground to b200 mesh and reacted with 100% phosphoric acid at 25 °C for 3 h in the case of calcite and at 25 °C for 3 days in the case of dolomite and ankerite. We extracted carbon dioxide from the carbonate samples according to the method of Al-Aasm et al. (1990) using a Finnigan Mat 251 mass spectrometer. The reproducibility of the analytical procedure was better than ±0.1‰ for calcite and ±0.2‰ for carbonate–Fe–Mg. All the samples were compared to a carbon dioxide reference obtained from a calcite standard prepared at the same time. Thus, oxygen isotope ratios were recalculated taking into account the fractionation factor for acid decomposition at 50 °C (1.01057 for ankerite and dolomite) and at 25 °C (1.01044 for calcite). C and O isotope measurements of carbonate materials of the Gulf of Cadiz have been previously carried out with this method (Díaz-del-Río et al., 2003; González et al., 2009). 2.4. Organic compound analysis (lipids and amino acids) To remove contaminants, we processed and cleaned the samples in parallel with procedural blanks and we eliminated the external surface of the carbonate pipes. We placed the powdered samples into distilled water (Milli-Q) and left them for 2 h in an ultrasonic bath at room temperature. Then we concentrated and dried the mixture under vacuum. Subsequently, we powdered the samples at 600 rpm using a planetary micro mill (Pulverisette 9, Fritsch, IdarOberstein, Germany). Since our samples were extensively altered, to characterize the preserved organic compounds we analyzed the samples by solid phase microextraction (SPME) coupled with gas chromatography– mass spectrometry (GC–MS), using the procedure of Menor-Salván

Table 1 Sampling site data for the carbonate pipes studied in this work. Dredge Anastasya Volcanic or no./cruise cruise physiographic unit 15

2001

2000

1, 2

2001

DI ¼ ∑ i

vari −AVG vari ·fac:coef :i STD vari

where vari is the original molar percent of each amino acid i, AVGvari and STDvari are the mean and standard deviations and fac.coef.i is the factor coefﬁcient of the ﬁrst axis of the PCA of Dauwe et al. (1999). The DI indicates the cumulative deviation with respect to an assumed average molar composition, with negative values indicating more degradation than the average and positive values indicating less. The DI varies from + 1.5 for labile, fresh material to − 3 for refractory, aged organic matter (Dauwe et al., 1999). 3. Results

Geographic Depth (m) Number of location samples studied

Coruña mud-volcano 36° 11′ N 7° 32′ W Arcos mud-volcano 36° 09′ N 7° 33′ W Iberico mound 36° 8′ N 7° 43′ W Cornide mound 36° 7′ N 7° 37′ W

et al. (2008). GC–MS analyses were carried out on an Autosystem XL-Turbo Mass Gold (Perkin Elmer) with an Elite-5 column (crossbond 5% diphenyl-95% dimethyl polysiloxane, 30 m × 0.25 mm i.d. × 0.25 μm film thickness) using He as carrier gas. The mass spectrometer was operated with the following parameters: mode, EI +; ionization energy, 70 eV; m/z range, 30–600; transfer line temperature, 300 °C. The temperature was programmed as follows: remain at 40 °C for 4 min, increase from 40 to 150 °C at 15 °C/min, remain at 150 °C for 2 min, increase from 150 to 255 °C at 5 °C/min, remain at 255 °C for 15 min, increase from 255 to 300 °C at 10 °C/min, and remain at 300 °C for 1 min. In order to detect long-chain hydrocarbons and hopanes, we subjected the samples (20 g) to liquid–solid extraction with 100 ml of nhexane:acetone (1:1) and three rounds of sonication. Combined extracts (300 ml) were subjected to the following steps: (a) filtration through a column with anhydrous sodium sulfate and freshly precipitated copper, (b) storage in hexane and concentration to 1 ml, (c) filtration through a cleaned glass-fiber microfilter, and (d) concentration to 25 μl under N2. We analyzed the extract (1 μl) by GC–MS as described above. For the molecular characterization of analytes, we used the reference spectra included in the NIST library and authenticated samples acquired from Sigma-Aldrich and Fluka. We considered an organic compound to be positively determined only when the correlation between the sample and reference spectra exceeded 85%. We determined amino acid concentrations by reverse-phase highpressure liquid chromatography (HPLC) according to Ruiz-Bermejo et al. (2007). Dissolved samples were hydrolyzed with 6 mol/l HCl at 110 °C for 24 h in sealed vials and then freeze-dried to remove water and HCl. We identified amino acids by comparing their retention times and UV absorption spectra with those of standards (Amino Acid Standards H, Pierce Chemical). We used the molar percent of the amino acids to calculate the degradation index (DI), developed by Dauwe and Middelburg (1998) and Dauwe et al. (1999). The DI assesses the diagenetic alteration of a sample by comparing it to a set of 28 samples in different degradation states and from different environments. We standardized the molar percent of individual amino acids by subtracting the mean of all values from individual results and dividing by the standard deviation of all measurements. The DI integrates the amino acid data weighted by the factor coefficients of the first axis of the principal component analysis (PCA) of Dauwe et al. (1999) according to the formula:

814

880

870–950

920–1145

3.1. Mineralogy, carbonate microfabrics and stable C and O isotopes The pipes showed a wide variety of shapes and ranged in size from several centimeters to a few decimeters. The external surface of the dry samples was brownish-gray and was densely perforated by small holes with diameters of 0.5–2 mm (Fig. 2A and B). These holes may be the result of chemical dissolution during exposure of the pipes to oxygenated waters. We observed grayish patches of seaﬂoor mud adhering to the walls and biological colonization by

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Fig. 2. Photographs of cylindrical carbonate pipes from the Gulf of Cadiz. (A) and (B) Pipes with the external surface densely perforated by small holes. (C) Transverse slab showing internal carbonate structure. (D) and (E) Pipes showing grayish patches of mud and biological colonization on their external walls.

incrusting organisms (serpulid worm tubes and small corals) over the external walls of the pipes (Fig. 2D and E). Internal carbonate was lighter and less porous, containing quartz, feldspar and phyllosilicate grains, and well-preserved remains of foraminiferal and ostracodal shells (Figs. 2C and 3). As shown previously, iron sulﬁdes and pseudomorph oxyhydroxides were common components of the internal structure of the pipes, forming framboidal agglomerates (Merinero et al., 2008, 2009). XRD analyses revealed that the carbonate pipes consisted of Ferich dolomite, ankerite and Mg-rich calcite with an admixture of calcite, goethite and quartz. Concentrations of quartz and calcite were the highest for pipes from the Arcos mud-volcano. Although carbonate pipes presented relatively homogeneous internal textures (Fig. 3), petrologic observations showed several internal carbonate fabrics: (1) small, elliptical pellets (5–20 μm) with a regular outline and clear edges embedded in the microcrystalline carbonate matrix (Fig. 4A); (2) clots of irregular shape and size, cloudy internal texture

and indistinct margins, surrounded by microcrystalline carbonate (Fig. 4B); and (3) spheroidal nodules with rims of small framboidal iron minerals (diameter b 2 μm) forming concentric layers and presenting a texture different from that of the surrounding carbonate matrix (Fig. 4C and D). Carbon and oxygen isotopic analyses were performed on samples collected from the Arcos mud-volcano and from the Iberico and Cornide mounds. The carbonate pipes were substantially depleted in 13 C: δ 13C values varied from −9.24‰ to −38.36‰. Carbonate δ 18O values ranged from −0.99‰ to +6.65‰. Lower values of δ 13C and δ 18O were obtained for pipes from the Arcos mud-volcano, where calcite was present at higher concentrations. 3.2. Composition of lipids and amino acids All samples showed similar amino acid compositions (Fig. 5). Of the 15 amino acids identiﬁed, the dominant amino acid in all samples

Fig. 3. Thin-section photomicrographs showing the internal texture of the carbonate pipes: a microcrystalline carbonate matrix containing grains of quartz (Q), detrital iron-oxides (O), and foraminiferal shells (F) ﬁlled with framboidal clusters of iron-oxyhydroxide pseudomorphs after pyrite (FR). Limits (L) are drawn between textures (pipe wall and ﬁlled internal channel).

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chain fatty alcohols were found in all samples, from C10 to C18 nalkanols. 4. Discussion 4.1. Relation between carbonate pipes and AOM

Fig. 4. Optical microscope images showing carbonate microfabrics of the studied pipes. Scale bar: 100 μm. (A) Peloids embedded in microcrystalline carbonate matrix. (B) Clotted microcrystalline carbonates. (C) A spherical nodule of microcrystalline carbonate surrounded by framboidal iron minerals.

was alanine, with a molar percent (mol%) varying from 36.5 to 44.9, followed by glycine (18.5–20.5 mol%), proline (3.3–10 mol%), histidine (2.3–6.4 mol%) and tyrosine (2.1–6.1 mol%). The remaining amino acids identiﬁed showed relative concentrations lower than 5 mol%. DI values were similar for all samples, ranging from − 0.43 to − 0.48. The lipids identiﬁed included C10 to C24 n-alkanes without predominance of odd-over-even carbon number, with a high abundance of compounds with 17–20 carbons (Fig. 6). Classical liquid–solid extraction, together with the SPME technique, showed a lack of hydrocarbons with more than 27 carbons. Other compounds detected were as follows: clusters of monomethyl-alkanes (MMA) and dimethylalkanes (DMA), phytanol (3,7,11,15-tetramethylhexadecanol), phytol (15-tetramethylhexadec-2-enol), and the regular isoprenoids farnesane (2,6,10-trimethyldodecane), nor-pristane (2,6,10-trimethylpentadecane), and pristane (2,6,10,14-tetramethylpentadecane). All samples contained appreciable amounts of the head-to-tail isoprenoid phytane (2,6,10,14-tetramethylhexadecane), which partially co-eluted with a minor amount of the tail-to-tail irregular C20-isoprenoid crocetane (2,6,11,15-tetramethylhexadecane). Linear and methyl-branched mid-

The observed depletion of 13C in the bulk carbonate is the first indication of the role played by the oxidation of methane in the formation of the carbonate pipes studied: δ 13C values varied from − 9.24‰ to −38.36‰. These δ 13C values are similar to those of other carbonates from the Gulf of Cadiz (Díaz-del-Río et al., 2003; Stadnitskaia et al., 2008). Depletion of 13C in carbonates is widely accepted evidence of a methane-related origin (Peckmann and Thiel, 2004). Methane oxidation results in enhanced concentrations of isotopically depleted bicarbonate in pore waters, resulting in carbonate precipitation (Ritger et al., 1987; von Rad et al., 1996; Peckmann and Goedert, 2005). These carbonates inherit the stable isotope composition of their carbon sources (Campbell et al., 2002; Peckmann and Thiel, 2004). However, the 13C content results from mixing of different carbon sources during carbonate precipitation and the extent of mixing is hard to determine. Thus, an evaluation of fluid composition of δ13C carbonate values alone is problematic: carbon sources other than hydrocarbons are usually relatively enriched in 13C and, thus, modern hydrocarbon-seep carbonates generally show higher δ13C values than their hydrocarbon source (Peckmann and Thiel, 2004). Therefore, if a certain amount of mixing of different carbon pools is assumed, the lowest δ13C value from the carbonate pipes (as low as −38.36‰) could indicate biogenic methane as a carbon source. However, the overall δ13C values point to thermogenic methane as the main carbon source according to previously reported studies of methane origin in active mud-volcanoes of the Gulf of Cadiz (Stadniskaia et al., 2006; Nuzzo et al., 2009). The high δ 18O value observed could be attributed to dolomite precipitation from 18O-enriched diagenetic fluids, rather than to a temperature effect. Precipitation of calcite in equilibrium with marine bottom water (δ 18O water ≈ 0‰ ± 0.2‰; Standard Mean Ocean Water) at 4 °C has a δ 18O value of approximately +3.2‰. Dolomite should be about 3‰ higher in δ 18O than coexisting calcite, and Mgcalcite should have a δ 18O value intermediate between these two phases at a given temperature (Anderson and Arthur, 1983; Land, 1983). Calcite precipitated from fresh water or at elevated temperatures should have significantly lower isotopic values. Pore fluids derived from destabilization of clathrate or clay mineral dehydration could cause a small increase (1–2‰) in the δ 18O of the local pore fluids (Martin et al., 1996). Both processes occur in different areas of the Gulf of Cadiz (Somoza et al., 2003; León et al., 2006; Hensen et al., 2007) and could therefore be considered likely fluid sources for precipitation of carbonate pipes. In addition, the lower values of δ 18O in samples from the Arcos mud-volcano may be explained by their higher content of calcite compared to other pipes in this study. Precipitation of carbonates and iron sulfides is the main mineralogical consequence of AOM coupled with sulfate reduction at methaneseep sites (Peckmann and Thiel, 2004). Vasconcelos et al. (1995) and Warthmann et al. (2000) showed that, under anoxic conditions, bacterial sulfate reduction (BSR) overcomes the kinetic barrier to dolomite formation by increasing the alkalinity of the carbonate and the surrounding medium. As discussed above, the lower values of δ13C indicate that a significant proportion of carbon in the bulk carbonate derived from methane. At the same time, the previously reported abundance of pyrite framboids in carbonate pipes of the Gulf of Cadiz (Merinero et al., 2008) indicates high local rates of sulfate reduction. Dolomite precipitation is enhanced by the removal of seawater sulfate from porewater, with concomitant increases in carbonate alkalinity (Baker and Kastner, 1981). The most plausible conclusion from these findings is that the dolomitic and ankeritic pipes in this study formed from pore fluids in which the seawater sulfate had been locally consumed.

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Fig. 5. Molar percents of the amino acids detected by HPLC in carbonate pipes from the Iberico and Cornide mounds and the Arcos mud-volcano. Abbreviations for amino acids are as follows: alanine (ALA), glycine (GLY), proline (PRO), histidine (HIS), tyrosine (TYR), aspartic acid (ASP), serine (SER), leucine (LEU), isoleucine (ILE), threonine, (THR), cysteine (CYS), glutamic acid (GLU), valine (VAL), phenylalanine (PHE), lysine (LYS).

The observed carbonate microfabrics provide additional evidence for the proposed relation between carbonate pipe formation and AOM. These features are frequently found in methane-related carbonates (e.g. Commeau et al., 1987; Hovland et al., 1987; Mazzini et al., 2004; Peckmann and Thiel, 2004; Han et al., 2008). Clotted microfabrics and carbonate nodules are considered signs of microbial activity (Burne and Moore, 1987; Coleman, 1993; Raiswell and Fisher, 2000;

Peckmann et al., 2003), and their occurrence may be related to small variations in the chemical environment during carbonate precipitation, caused by metabolic activity of methane-oxidizing organisms (Peckmann et al., 2002; Peckmann and Thiel, 2004; Buggisch and Krumm, 2005). Pellets, for their part, could be interpreted as a product of bacterially induced precipitation of carbonates (Chavetz, 1986).

Fig. 6. Typical lipid distribution patterns in the carbonate pipes studied from the Gulf of Cadiz. (A) GC–MS ion chromatogram (m/z = 85) where numbers denote the number of carbons of n-alkanes, Fa = farnesane, NP = nor-pristane, Pr = pristane, Ph = phytane, Cr = crocetane, stars = unidentiﬁed methyl-alkanes. (B) GC–MS ion chromatogram (m/z = 184) showing phytanol, phytol and fatty alcohol peaks.

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4.2. Aerobic alteration of carbonate pipes Before their collection in this study, the carbonate pipes sampled in the Gulf of Cadiz lay on the seafloor and were in contact with oxygenated water (Díaz-del-Río et al., 2003), so aerobic alteration may be expected. Indeed, the pipes presented visible signs of bioerosion and alteration: the color and the small holes perforating the external surface may reflect dissolution due to the acidity generated by iron sulfide oxidation during the exposure of the pipes to oxygenated waters. Iron sulfide oxidation and pseudomorphic transformation into iron oxyhydroxides have been documented in previous mineralogical studies on these pipes (Merinero et al., 2008, 2009). Colonization by incrusting organisms and grayish patches of mud attached to the walls of the pipes are also signs of their residence on the seafloor and subsequent exposure to oxygenated waters. This hypothesis of aerobic alteration of the pipes is supported by DI values calculated from the molar percent of the amino acids. Values ranged from −0.43 to − 0.48, indicating degradation of the organic matter preserved within the pipes, based on the DI interpretation of Dauwe et al. (1999). Alanine was the most abundant amino acid identified, followed by glycine. Both alanine and glycine are preferentially preserved in degraded organic matter, and their relative concentration increases with degradation (Lee and Cronin, 1984; Müller et al., 1986; Lee et al., 2000). On the other hand, valine, isoleucine, leucine and histidine were found at low mol% in the pipes. These amino acids tend to be preferentially decomposed during organic matter degradation (Hecky et al., 1973). 4.3. Sources of organic compounds The inventory of organic substances preserved in methane-derived carbonates reflects the combination of allochthonous organic matter preserved in carbonate porosity (e.g. filling foraminiferal tests), organic compounds derived from methanotrophic and sulfate-reducing organisms and seeped hydrocarbons trapped within carbonates. In addition, biodegradation, thermal maturation, and aerobic oxidation influence the nature of the organic substances preserved inside these carbonates, altering the composition and diluting their concentrations. These processes alter the preserved organic substances and their concentrations to different extents over time and must be taken into account when interpreting the patterns observed in methane-derived carbonates. Pristane, phytane, crocetane, farnesane and nor-pristane were the isoprenoids detected within carbonate pipes. Crocetane was the only specific biomarker for AOM identified, and its precise contents are difficult to establish since the abundance of phytane partially obscured its presence. Studies of crocetane in fossil samples have shown that this compound can be used as a diagenetically stable fingerprint for methanotrophic archaea that thrive in methane-rich settings and that aid in the formation of methane-derived carbonates (Peckmann and Thiel, 2004). The structure of crocetane, involving a tail-to-tail linkage of isoprene units, suggests an archaeal origin (Elvert et al., 1999, 2000; Peckmann et al., 1999; Thiel et al., 1999, 2001; Pancost et al., 2000). However, the association between crocetane and methanotrophic archaea cannot be definitively demonstrated without determining δ13C signatures. Pristane and phytane are typical diagenetic products, respectively, of algal tocopherols and the phytol side-chain of chlorophyll (Didyk et al., 1978; Goossens et al., 1984). However, some authors, such as Birgel et al. (2006), have proposed that sedimentary phytane, pristane, nor-pristane and farnesane can be derived from lipids of methane-oxidizing archaea. In particular, the most plausible biological precursors of phytane are archaeol and hydroxyarchaeol derivatives (Peckmann and Thiel, 2004). Similar scenarios may be valid for pristane, nor-pristane and farnesane, which may derive from archaeal isopranyl lipids (Peckmann et al., 2002; Goedert et al., 2003). Therefore, we speculate that the observed pattern of isoprenoids is derived from aerobic degradation of organic compounds

from sulfate dependent AOM with better preservation of archaeal than bacterial lipids. This is because the carbon skeletons of lipids from SRB, due to their structure and low molecular weight, are less resistant to biodegradation, thermal alteration, and weathering than are isoprene-based archaeal biomarkers (Peters and Moldowan, 1993). Consequently, these processes may erase evidence of an SRB source faster than evidence of archaeal chemofossils in methanederived carbonates. Other organic compounds detected inside carbonate pipes were phytanol, phytol and clusters of monomethyl-alkanes (MMA) and dimethyl-alkanes (DMA). Phytanol, a possible precursor of pristane (Peters et al., 2005), is a compound frequently found in AOM environments of putative archaeal origin (Thiel et al., 1999; Hinrichs and Boetius, 2002; Peckmann et al., 2004). Oxic conditions promote the conversion of phytanol to pristane by oxidation of phytanol to phytanic acid, decarboxylation to pristene, and then reduction to pristane. Therefore, the detection of pristane in our organic analyses may be explained by the exposure of the carbonate pipes to oxygenated waters and the subsequent degradation of some precursor formed during AOM and initially trapped in the carbonate matrix. Detection of phytol could be interpreted as allochthonous input because it is frequently derived from the ester-linked phytol moiety in chlorophyll and hence is associated with photosynthetic activity of algae and higher plants. The branched hydrocarbons methylalkanes (MMAs) and dimethylalkanes (DMAs) are recognizable in ancient sediments. They are diagenetic products derived from biosynthesized functionalized lipid precursors (Summons, 1987; Summons et al., 1988) and can be considered potential biomarkers. Their use as specific, reliable biomarkers is limited by inadequate knowledge about the ranges in which they occur in modern sedimentary environments and by the likelihood that they serve as precursors in multiple biogenic pathways. Nevertheless, since they are precursors of monomethyl fatty acids in some common SRB (Dowling et al., 1986), we speculate that a possible source of the short-chain MMAs and DMAs detected in our study is the degradation of long-chain fatty acids synthesized by SRB initially present within the carbonate pipes. Finally, we found linear and methyl branched mid-chain fatty alcohols in all samples, from C10 to C18. Fatty alcohols frequently occur in sediments at modern seeps and their presence has been attributed to SRB performing AOM (Orphan et al., 2001). However, they cannot be used to determine the specific groups of microorganisms involved in the way that fatty acids can. 5. Summary and conclusions Carbonate pipes sampled from the seafloor of the Gulf of Cadiz are dominated by authigenic microcrystalline Fe-rich dolomite, ankerite and Mg-rich calcite. The δ 13C values of carbonate pipes indicate formation from oxidation of methane. Their δ 18O values agree with dolomite precipitation, which may have occurred in a regime of sulfate depletion compatible with BSR. The observed carbonate microfabrics of clotted carbonates and spherical nodules of microcrystalline carbonate surrounded by framboidal iron minerals reflect the activity of microorganisms that drive AOM. Despite the observed intensive aerobic alteration, some organic compounds are preserved in these carbonate pipes. Crocetane was the only lipid identified from methanotrophic archaea. Although the observed pattern of organic compounds (n-alkanes, regular isoprenoids and alcohols) indicates aerobic alteration of the pipes, they are also compatible with an origin in AOM. Additional signs of exposure to oxygenated waters and aerobic alteration are visible in the external surface of the pipes, such as small holes, grayish color, colonization and bioerosion. Moreover, the degradation indexes calculated from the molar percent of amino acids preserved in pipes also indicate aerobic degradation of the trapped organic compounds. Therefore, the conclusions of this study point to a significant microbial contribution to the precipitation of the methane-derived

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carbonate pipes sampled in the Gulf of Cadiz. We therefore speculate that the precipitation is due, at least in part, to AOM coupled with BSR. These findings are not definitive, since aerobic alteration masked the lipid biomarkers of the specific microbes involved. Acknowledgments This study forms part of the work of the Research Group CAM-UCM and the work on terrestrial analogs for the exploration of Mars at the Centro de Astrobiología. The results were obtained within the framework of the European Science Foundation EuroCORE-EuroMARGINS projects “MOUNDFORCE” (01-LEC-EMA06F, REN-2002-11668-E-MAR) and “MVSEIS” (01-LEC-EMA24F, REN-2002-11669-E-MAR). We thank all the scientific and technical personnel who participated in the oceanographic cruises Anastasya 2000 and 2001 of the Cornide de Saavedra research vessel. We are also very grateful to the “Centro de Microscopía Electrónica Luis Bru” (Complutense University of Madrid), to Drs. J.A. Martín-Rubí and F.J. González of the Geological Survey of Spain and to Dr. A. Delgado of the “Estación Experimental del Zaidín” of the CSIC (Granada, Spain) for their contributions to the study. We wish to thank the collaborative efforts and insightful remarks of Dr. Victor Díaz del Río (Centro Oceanográfico de Málaga, Instituto Español de Oceanografía). We appreciate the careful revision of Armando Chapin Rodríguez who greatly aided in improving the final manuscript. References Al-Aasm, I.S., Taylor, B.E., South, B., 1990. Stable isotope analysis of multiple carbonate samples using selective acid extraction. Chemical Geology: Isotope Geoscience Section 80, 119–125. Anderson, F.F., Arthur, M.A., 1983. Stable isotopes of oxygen and carbon and their application to sedimentologic and paleoenvironmental problems. In: Arthur, M.A., Anderson, T.F., Kaplan, I.R., Veizer, J., Land, L.S. (Eds.), Stable Isotopes in Sedimentary Geology 10. SEPM Short Course Notes, Chapter 1, pp. 1–151. 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The application of stable isotopes to studies of the origin of dolomite and to the problems of diagenesis of clastic sediment. In: Arthur, M.A., Anderson, T.F., Kaplan, I.R., Veizer, J., Land, L.S. (Eds.), Stable Isotopes in Sedimentary Geology 10, pp. 1–22. SEPM Short Course Notes, Chapter 4. Lee, C., Cronin, C., 1984. Particulate amino acids in the sea: effects on primary productivity and biological decomposition. Journal of Marine Research 42, 1075–1097. Lee, C., Wakeham, S.G., Hedges, J.I., 2000. Composition and flux of particulate amino acids and chloropigments in the equatorial Pacific seawater and sediments. Deep-Sea Research Part I: Oceanographic Research Papers 47, 1535–1568. León, R., Somoza, L., Medialdea, T., Maestro, A., Díaz-Del-Río, V., Fernandez-Puga, M.C., 2006. Classification of sea-floor features associated with methane seeps along the Gulf of Cádiz continental margin. Deep-Sea Research Part II 53, 1464–1481. León, R., Somoza, L., Medialdea, T., González, F.J., Díaz del Río, V., Fernández-Puga, M.C., Maestro, A., Mata, M.P., 2007. Sea-floor features related to hydrocarbon seeps in deepwater carbonate-mud mounds of the Gulf of Cádiz: from mud flows to carbonate precipitates. Geo-Marine Letters 27, 237–247. Maestro, A., Somoza, L., Medialdea, T., Talbot, C.J., Lowrie, A., Vázquez, J.T., Díaz-del-Río, V., 2003. Large-scale slope failure involving Triassic and Middle Miocene salt and shale in the Gulf of Cádiz (Atlantic Iberian Margin). Terra Nova 15, 380–391. Maldonado, A., Somoza, L., Pallarés, L., 1999. The Betic orogen and the Iberian-African boundary in the Gulf of Cádiz: geological evolution (central North Atlantic). Marine Geology 155, 9–43. Martin, J.B., Kastner, M., Henry, P., LePichon, X., Lallement, S., 1996. Chemical and isotopic evidence of sources of fluids in a mud volcano field seaward of the Barbados accretionary wedge. Journal of Geophysical Research 101, 20325–20345. Mazzini, A., Ivanov, M.K., Parnell, J., Stadnitskaia, A., Cronin, B.T., Poludetkina, E., Mazurenko, L., van-Weering, T.C.E., 2004. Methane-related authigenic carbonates from the Black Sea: geochemical characterisation and relation to seeping fluids. Marine Geology 212, 153–181. Medialdea, T., Vegas, R., Somoza, L., Vázquez, J.T., Maldonado, A., Díaz-del-Río, V., Maestro, A., Córdoba, D., Fernández-Puga, M.C., 2004. Structure and evolution of the “Olistostrome” complex of the Gibraltar Arc in the Gulf of Cádiz (eastern Central Atlantic): evidence from two long seismic cross-sections. Marine Geology 209, 173–198. Medialdea, T., Somoza, L., Pinheiro, L.M., Fernández-Puga, M.C., Vázquez, J.T., León, R., Ivanov, M.K., Magalhaes, V., Díaz-del-Río, V., Vegas, R., 2009. Tectonics and mud volcano development in the Gulf of Cádiz. Marine Geology 261, 48–63. Menor-Salván, C., Ruiz-Bermejo, M., Osuna-Esteban, S., Muñoz-Caro, G., VeintemillasVerdaguer, S., 2008. Synthesis of polycyclic aromatic hydrocarbons and acetylene polymers in ice: a prebiotic scenario. Chemistry and Biodiversity 5, 2729–2739. Merinero, R., Lunar, R., Martínez-Frías, J., Somoza, L., Díaz-del-Río, V., 2008. Iron minerals in hydrocarbon seeps related carbonates, Gulf of Cadiz (southwest Iberian Peninsula). Marine and Petroleum Geology 25, 706–713. Merinero, R., Lunar, R., Somoza, L., Díaz-del-Río, V., Martínez-Frías, J., 2009. Nucleation, growth and oxidation of framboidal pyrite associated with hydrocarbon-derived submarine chimneys: lessons learned from the Gulf of Cadiz. European Journal of Mineralogy 21, 947–961. Mulder, T., Lecroart, P., Hanquiez, V., Marches, E., Gonthier, E., Guedes, J.C., Thiébot, E., Jaaidi, B., Kenyon, N., Voisset, M., Perez, C., Sayago, M., Fuchey, Y., Bujan, S., 2006.

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The western part of the Gulf of Cadiz: contour currents and turbidity currents interactions. Geo-Marine Letters 26, 31–41. Müller, P.J., Suess, S., Ungerer, C.A., 1986. Amino acids and amino sugars of surface particulate and sediment trap material from waters of the Scotia Sea. Deep-Sea Research Part I: Oceanographic Research Papers 33, 819–838. Niemann, H., Duarte, J., Hensen, C., Omoregie, E., Magalhães, V.H., Elvert, M., Pinheiro, L.M., Kopf, A., Boetius, A., 2006. Microbial methane turnover at mud volcanoes of the Gulf of Cadiz. Geochimica et Cosmochimica Acta 70, 5336–5355. Nuzzo, M., Hornibrook, E.R.C., Gill, F., Hensen, C., Pancost, R.D., Haeckel, M., Reitz, A., Scholz, F., Magalhães, V.H., Brückmann, W., Pinheiro, L.M., 2009. Origin of light volatile hydrocarbon gases in mud volcano ﬂuids, Gulf of Cadiz — evidence for multiple sources and transport mechanisms in active sedimentary wedges. Chemical Geology 266, 350–363. Orphan, V.J., House, C.H., Hinrichs, K.-U., McKeegan, K.D., DeLong, E.F., 2001. Methaneconsuming archaea revealed by directly coupled isotopic and phylogenic analysis. Science 293, 484–487. Pancost, R.D., Damsté, J.S.S., deLint, S., van-der-Maarel, M.J.E.C., Gottschal, J.C., 2000. Biomarker evidence for widespread anaerobic methane oxidation in Mediterranean sediments by a consortium of methanogenic archaea and bacteria. Applied and Environmental Microbiology 66, 1126–1132. Paull, C.K., Chanton, J.P., Neumann, A.C., Coston, J.A., Martens, C.S., 1992. Indicators of methane-derived carbonates and chemosynthetic organic carbon deposits examples from the Florida escarpment. Palaios 7, 361–375. Peckmann, J., Goedert, J.L., 2005. Geobiology of ancient and modern methane-seeps. Palaeogeography, Palaeoclimatology, Palaeoecology 227, 1–5. Peckmann, J., Thiel, V., 2004. Carbon cycling at ancient methane-seeps. Chemical Geology 205, 443–467. Peckmann, J., Thiel, V., Michaelis, W., Clari, P., Gaillard, C., Martire, L., Reitner, L., 1999. Cold seep deposits of Beauvoisin (Oxfordian; southeastern France) and Marmorito (Miocene; northern Italy): microbially induced authigenic carbonates. International Journal of Earth Sciences 88, 60–75. Peckmann, J., Goedert, J.L., Thiel, V., Michaelis, W., Reitner, J., 2002. A comprehensive approach to the study of methane-seep deposits from the Lincoln Creek Formation, western Washington State, USA. Sedimentology 49, 855–873. Peckmann, J., Goedert, J.L., Heinrichs, T., Hoefs, J., Reitner, J., 2003. The Late Eocene ‘Whiskey Creek’ methane-seep deposit (western Washington State): part II. Petrology, stable isotopes, and biogeochemistry. Facies 48, 241–254. Peckmann, J., Thiel, V., Reitner, J., Taviani, M., Aharon, P., Michaelis, W., 2004. A microbial mat of a large sulfur bacterium preserved in a Miocene methane-seep limestone. Geomicrobiology Journal 21, 247–255. Peters, K.E., Moldowan, J.M., 1993. The Biomarker Guide—Interpreting Molecular Fossils in Petroleum and Ancient Sediments. Prentice-Hall, Englewood Cliffs, New Jersey. 363 pp. Peters, K.E., Walters, C.C., Moldowan, J.M., 2005. The Biomarker Guide, second edition. Cambridge University Press, United Kingdom. 490 pp. Pinheiro, L., Ivanov, M.K., Sautkin, A., Akhmanov, G., Magalhães, V., Volkonskaya, A., Monteiro, J.H., Somoza, L., Gardner, J., Hamouni, N., Cunha, M.R., 2003. Mud volcanism in the Gulf of Cádiz: results from the TTR-10 cruise. Marine Geology 195, 131–151.

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TECHNICAL NOTE

www.rsc.org/jem | Journal of Environmental Monitoring

Monitoring the fall of large atmospheric ice conglomerations: a multianalytical approach to the study of the Mejorada del Campo megacryometeor Francisco Alamilla Orellana,a Jose´ Ma Ramiro Alegre,a Jose´ Carlos Cordero Pe´rez,a Ma Paz Martıń Redondo,b Antonio Delgado Huertas,c Ma Teresa Fernańdez Sampedro,b Ce´sar Menor-Salvań,b Marta Ruiz-Bermejo,b Fernando Lo´pez-Vera,d Jose´ A. Rodrı´guez-Losadae and Jesus Martinez-Frias*b Received 5th December 2007, Accepted 29th January 2008 First published as an Advance Article on the web 25th February 2008 DOI: 10.1039/b718785h Certain local atmospheric anomalies, such as the formation of unusually large ice conglomerations (megacryometeors), have been proposed to be a potential natural hazard for people and aviation, as well as geoindicators for fingerprinting larger-scale atmospheric environmental changes. On March 13th 2007, at approximately 10:15 am, an ice chunk weighing about 10 kg fell from the clear-sky and crashed through the roof (around 15 m) of an industrial storage house in Mejorada del Campo, a town located 20 km east from Madrid. The megacryometeor monitoring follow-up and the original investigation presented here includes, for the first time, both logistic and scientific collaboration between the Laboratory of the Environment, Criminalistic Service (SECRIM, the Spanish ‘‘Guardia Civil’’) and academic and scientific institutions (universities and the Spanish National Research Council). We propose that the management procedure of the incident, along with the detailed scientific research and combination of analytical methodologies in different laboratories, can serve as a protocol model for other similar events.

Introduction Environmental monitoring is the foundation for selecting management approaches and safety procedures, for developing predictive modelling and process research, and for integrating the scientific information necessary to make key decisions.1 In accordance with current scientific priorities for the International Global Atmospheric Chemistry (IGAC),2 the new directions in environmental strategies encourage the study of the atmosphere in its entirety, taking into account the interactions and modifications occurring at different scales. Certain local atmospheric anomalies, such as the formation of unusually large ice conglomerations (megacryometeors), have been proposed to be a potential natural hazard3 for people and aviation, as well as geoindicators4 for fingerprinting larger-scale atmospheric chemical and physical changes.5–12 By the study of atmospheric soundings from NOAA, and NCEP/NCAR reanalysis data of the upper troposphere, the formation of megacryometeors has been linked to undulations of the tropopause (mean upper tropospheric temperature gradient for 19 ice fall events occurred

a Direccioń General de la Guardia Civil, Servicio de Criminalı´stica, Laboratorio de Medioambiente, C\Guzmań el Bueno 110, Madrid, Spain b Centro de Astrobiologia, CSIC/INTA, Asociado al NASA Astrobiology Institute, Ctra de Ajalvir, km. 4, Torrejoń de Ardoz, Madrid, 28850, Spain. E-mail: martinezfj@inta.es; Fax: +34-91-5201621 c Departmento de Ciencias de la Tierra y Quı´mica Ambiental, Estacioń Experimental del Zaidıń, CSIC, Prof. Albareda 1, Granada, 18008, Spain d Departamento de Geologıá y Geoquı´mica, Facultad de Ciencias, Universidad Autońoma de Madrid, Madrid, 28049, Spain e Departamento de Edafologıá y Geologıá, Facultad de Biologıá, Universidad de La Laguna, La Laguna, Tenerife, C. Islands, 38206, Spain

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for the past five years combined was significantly greater than climate normals), ozone anomalies and strong wind turbulence.6,9,11,12 This is of relevance from the environmental point of view, given that observations suggest that: (1) the mixing ratio of water vapour in the stratosphere has increased by 20–50% from the 1960s to the mid-1990s13 and (2) interchanges of water vapour, favoured by tropopause disturbances, can play a central role in atmospheric chemistry, influencing heterogeneous chemical reactions, with subsequent implications in climate change, as part of a global long-term trend. Cooler stratospheric temperatures appear when there is more water vapor present, and water vapor also leads to the breakdown of ozone molecules.14,15 The megacryometeor monitoring follow-up and the original investigation presented here on the most recent ice fall event, which occurred in Spain in 2007, includes for the first time both logistic and scientific collaboration between the Laboratory of the Environment, Criminalistic Service (SECRIM, Spanish ‘‘Guardia Civil’’) and academic and scientific institutions (universities and the Spanish National Research Council). We propose that the management procedure of the incident, along with the scientific research and combination and comparison of analytical methodologies in different laboratories, can serve as a protocol model for other similar events.

Monitoring and research of megacryometeor incidents Our monitoring and research study of the clear-sky ice fall events was initiated in January 2000, in the context of a multidisciplinary scientific commission that was coordinated by the Spanish National Research Council (CSIC). A specific website was established, electronically hosted by the Thematic Network This journal is ª The Royal Society of Chemistry 2008

of Earth Sciences in Spain ‘Tierra’, in which pictures, incident information and scientific results were included.16 After almost eight years of monitoring and research, it has become evident that megacryometeors are not classical big hailstones, ice from aircrafts (waste water or tank leakage), nor the simple result of icing processes at high altitudes. They display textures, hydrochemical features and oxygen and hydrogen isotopic values which unequivocally confirm they are the result of complex formation processes within the atmosphere. More specifically, the megacryometeors’ water is consistent with the Craig’s Global Meteoric Water Line17 and their isotopic ranges ( 25& > dDSMOW > 127&; 4.52& > d18OSMOW > 17.25&) are clearly tropospheric.11 At present, no model is able to satisfactorily explain what factors cause the ice nucleation and growth,18,19 or how megacryometeors can actually be formed and maintained in the atmosphere. Nevertheless, any model should involve the existence of an ice-supersaturated region (ISSR), that is, a supersaturated but cloud free airmass,20,21 connected with extreme atmospheric turbulence, associated with the observed tropopause undulation.6,11 In this sense, theoretical calculations, based on experimentally-obtained dDSMOW variations, indicate that the vertical trajectory in effective growth of the megacryometeors was lower than 3.2 km.11 It is important to note that a detailed historical review of such ice fall events verifies that there are many documented references of similar falls of large blocks of ice which go back to the first half of the 19th century (prior to the invention of aircrafts).22–24 It is still too early to ascertain whether there is a real multiplication effect of the number of megacryometeor incident exclusively due to natural causes (now there is a statistical artifact in the analysis as the information circulates very fast and we can know rapidly what is happening in different parts of the world). Nevertheless, and mainly after 1950, the number of megacryometeor hits has apparently increased. More than 100 events have been witnessed and recorded, affecting practically the entire planet (Argentina, Australia, Austria, Canada, Colombia, India, Italy, Japan, Mexico, New Zealand, Portugal, South Africa, Spain, Sweden, The Netherlands, United Kingdom and the USA).3,11,12 From 2001 to 2006, a total of 52 ice-fall events have been witnessed and recorded. Verifiable effects include the megacryometeors’ crashing (some of them weighing more than 100 kg) through roofs or producing small impact craters (i.e., La Milana, Soria, Spain; Surrey, UK; Oakland, California, USA).25 As of this writing, 12 new documented ice falls have been recorded in 2007, eight in the USA, two in The Netherlands, one in the UK and one in Spain: the Mejorada del Campo megacryometeor.

Incident description On March 13th 2007, at around 10:15 am, an ice chunk weighing roughly 10 kg fell from clear-sky and crashed through the roof (around 15 m) of an industrial storage house in Mejorada del Campo, a town located 20 km east from Madrid.26 The workers were in the interior of the premises, and were witnesses of the ‘‘dry and very strong noise’’ caused by the impact. Due to the arrangement and elongated shape of the metallic material plates of the roof, a pronounced dent and an irregular hole of around 2 m x 1 m, were produced. A significant part of the ice chunk stayed deposited on the concavity of the dented part of the This journal is ª The Royal Society of Chemistry 2008

Fig. 1 Megacryometeor which fell on 13th March 2007, crashing through the roof (around 15 m) of an industrial storage house in Mejorada del Campo, Madrid. The picture shows the fragments of the ice chunk, which stayed deposited in the concavity of the dented part of the roof (see text for approximate weight and size).

roof (Fig. 1), whereas other ice fragments fell into the interior of the industrial storage room. The plates of plaster underlying the metallic material of the roof were also broken by the impact. The ice was white and semi-transparent, displaying a nearly equidimensional arrangement of the fragments (Fig. 1). Fortunately, the incident caused only material damage and nobody was hurt. Experts from the SECRIM, investigated the incident, ruling out other hypotheses (e.g. vandalism). Some ice pieces were collected and transported using portable freezers (the same day of the event), to perform the first set of hydrochemical analyses in the SECRIM’s Laboratory of Environment of the ‘‘Guardia Civil’’, and the rest of the ice was preserved under frozen conditions for later studies. Subsequently, a second sampling of the ice was carried out on April 13th 2007. Various megacryometeor fragments were moved, also under controlled conditions, by two SECRIM members (one of them is the first author of the present article), to the Planetary Geology Laboratory of the Centro de Astrobiologı´a (CAB).

Experimental The ice samples (several pieces weighing a total of 842 g) were kept in aseptic bags and immediately stored under refrigeration at approximately 20 C, to avoid textural changes, as well as to prevent possible contamination on the megacryometeor surface by water-steam condensation, or by the absorption of carbon dioxide from the environment. A previous stage of the characterization analysis was to remove the external part of the ice using an aseptic cutter. The set of analyses performed (all of them from liquid aliquots of the ice samples) comprises the combination of pH and conductivity, differential scanning calorimetry (DSC), ion chromatography (IC), inductively coupled plasma mass spectroscopy (ICP-MS), electrothermal atomic absorption spectrometry (ETAAS) with a graphite furnace, stable isotope mass spectrometry (SIMS), solid phase micro-extraction-gas chromatography-mass spectrometry (SPME-GC-MS), and microbiological analysis. J. Environ. Monit., 2008, 10, 570–574 | 571

The pH and conductivity determination was carried out using a CRISON pH-meter (mod. Basic 20) and a MeterLab PHM220 pH-meter, and a CDM210 conductivity-meter (Radiometer, Copenhagen, Denmark), respectively (NIST standard calibration). Independent measurement analyses were performed in SECRIM and CAB laboratories. Thermal analysis of selected ice samples was performed in a DSC 2920 of TA Instruments. Temperature and heat flow were calibrated in the common manner, using the onset temperatures of the melting of indium (429.75 K) and an empty pan was used as reference. The sample (20–25 mg) was put into a hermetically sealed aluminium pinhole pan. The sample was cooled from ambient down to 50 C, and reheated at 3 C min 1 up to 150 C. For a comparison, this same procedure was also applied to MilliQ water and tap water from the locality of Mejorada del Campo. The anion content was determined using a Dionex LC20 chromatograph equipped with a ED 40 electrochemical detector, a GP50 gradient pump and a AS 40 autosampler, a CD25 suppressed conductivity detector and an anion selfregenerating suppressor ASRS ULTRA II, 4 mm (Autosuppresion Recycle mode). Quantitative multielemental analysis was performed by means of an inductively coupled plasma source mass spectrometer (ELAN9000 Q-ICP-MS) equipped with a Ryton cross-flow nebulizer, a Scott spray chamber and a Cetac ASX-510 autosampler. The sampler transport to the nebulizer was established by a peristaltic pump. In order to obtain maximum precision, the instrument was optimized daily, using a solution containing 10 ppb of Mg, Cu, Rh, Cd, In, Ba, Ce, Pb and U to obtain maximum 103Rh intensity, as well as an oxide and double charge ion levels (lower than 3%). Regarding the reagents and standard samples, multielement external standard working solutions were prepared by accurate dilution of two commercial ICP-MS standards, covering the whole mass range: Certipur ICP multielement standard solution VI (Merk) and multielement calibration solution 2 (PerkinElmer). The acids employed in the sample treatment are suprapur grade (Merk), and high purity water (18.2 mU) from a MilliQ water system (Millipore) was used. Several aliquots were sampled from the interior of the megacryometeor. Likewise, additional tap water samples from Mejorada del Campo and Madrid, and Madrid rainwater were analyzed for comparison following the same analytical routine. The quality control of the analysis process was studied monitoring the recovery of the internal standard during the analysis and of all the elements in the quality control standard. Determination of arsenic content was specifically performed by a Perkin Elmer AAnalyst 600 atomic absorption spectrometer, equipped with a longitudinal Zeeman-effect background corrector and an AS-800 autosampler. The standard PerkinElmer THGA transversely heated graphite furnace atomizer with integrated platform was used. An electrode-less discharge lamp (EDL) was used for the determination of As (l ¼ 193.7 nm). The isotopic study was carried out at the Stable Isotope Laboratory of the Estacio´n Experimental del Zaidı´n (Granada, Spain). Oxygen in water was analysed by the CO2–H2O equilibration method.27,28 To determine hydrogen isotopic ratios we used reduction with Zn at 450 C.29,30 Isotopic ratios were measured by a Finnigan MAT 251 mass spectrometer. The experimental error was 572 | J. Environ. Monit., 2008, 10, 570–574

0.1& and 1& for oxygen and hydrogen, respectively, using EEZ-3 and EEZ-4 as internal standards that were previously calibrated vs. V-SMOW, SLAP and GIPS water. An organic compounds analysis was performed by solid phase microextraction (SPME) coupled with GC-MS: c.a. 4 mL of the megacryometeor sample was heated in a vial closed with a septum at 70 for 45 min. A 100 mm CAR-polydimethylsiloxane (CARPDMS) fibre was then exposed to the headspace, keeping the sample at the same temperature for a further 45 min. Analytes on the fibre were then thermally desorbed in the injection port of a Perkin Elmer Autosystem XL-Turbomass GC-MS instrument at 290 for 4 min (split less mode). The analysis was performed using a capillary column (5% diphenyl–95% dimethylpolysiloxane, 30 m 0.25 mm ID, 0.25 mm film) and using He as carrier gas. The temperature was raised from 40 (4 min) to 150 at a rate of 15 min 1, held for 2 min, 150 to 255 at 5 min 1, held for 15 min, and 255 to 300 at 10 min 1, and held for 1 min. The mass spectrometer was operated under EI mode, at an ionization energy of 70 eV, m/z range 30–600, transfer line at 300 . In addition to all these techniques, a microbiological analysis of the external part of the ice samples was also carried out.

Results and discussion The megacryometeor water has a pH in the range of from 7.05 to 7.86 ( 0.10), and conductivity values from 56.4 to 69.2 ( 7) mS cm 1. Thermal analysis of the ice indicates melting values ranging from 0.09 C to 3.12 C (316.4 J g 1) and boiling values from 99.34 C–104.32 C (1959 J g 1) (Fig. 2). Likewise, main ranges (mg l 1) of F, BrO3, Cl, NO2, NO3, PO4 and SO4 are the following: F: 0.52–0.68; BrO3: <LQ 0.13; Cl: 6.41–8.37; NO2: <LQ; PO4: 0.70–0.83 and SO4: 3.38–3.73. Table 1 displays a summary of the main hydrochemical results of the ice obtained by ICP-MS. Arsenic content was below detection limit (1.08 ppb) in all cases. Isotopically, the distribution of the samples (35 points of isotopic analyses covering different parts of the ice fragments; 9.76& > d18OSMOW > 10.70& and 49& > dDSMOW > 56&) in the Craig’s line verify that they match the Meteoric Water Line,17 providing evidence of a direct condensation of the ice from atmospheric (unequivocally tropospheric) water

Fig. 2 DSC of the ice, showing the ranges of melting and boiling values.

This journal is ª The Royal Society of Chemistry 2008

Table 1 ICP-MS hydrochemical results showing the ice composition of the Mejorada del Campo megacryometeor (MJC-m). MJC-tw: Mejorada del Campo tap water. M-rw: Madrid rainwater. M-tw: Madrid tap water. BDL: below detection limit. ND: non detected, LOD: limit of detection. The operating parameter setting is the following: RF power (W): 1000. Nebulizer gas flow: 0.89 l min 1. Lens voltage: 7.25 volts. Analog Stage voltage: 1900 volts. Pulse stage voltage: 1100 volts. Sweep/reading: 6. Reading/replicate: 1. Replicate: 1 ppb

MJC-m

MJC-tw

M-rw

M-tw

LOD

Ca Na Mg Al Si P K Sc V Cr Mn Fe Co Ni Cu Zn Ge As Rb Sr Y Zr Cd Ba La Ce Pr Nd Sm Pb

5333.6 7900.3 655.1 130.3 253.3 239.7 1853.6 BDL BDL 2 10 74.8 BDL 2.239 36.5 180.2 ND BDL BDL 26 0.1 ND 2.1 12.2 BDL BDL BDL BDL BDL 4

9430.1 5931.1 1705 76.5 2089.3 ND 792.9 1 BDL 1.7 1.7 30.3 BDL BDL 3.6 4.6 ND BDL BDL 44.7 BDL ND BDL 6.9 BDL BDL BDL BDL BDL BDL

7437.8 1342 376.8 80.1 280.1 110.5 603.1 BDL 1.9 ND 12.3 39.1 0.7 1.6 11.7 88.2 ND BDL BDL 23.3 0.1 BDL BDL 11.1 BDL BDL BDL BDL BDL 5.5

7655.5 4494.9 1308.3 388.7 1766.7 5.1 610.5 0.9 BDL 1.2 9.6 268.4 BDL 1.6 491.1 136.7 ND BDL BDL 31.1 0.1 ND ND 6.8 BDL BDL BDL BDL BDL 0.5

3.79 1.25 0.81 0.51 1.29 2.86 6.49 0.17 1.02 0.32 0.23 2.73 0.4 0.9 1.29 0.84 0.75 1.08 3.22 2.37 0.1 0.25 0.85 2.42 1.74 1.38 1.63 0.25 0.37 0.25

vapour. Regarding the analysis by GC-MS, it is important to note that no organic compounds were detected in the ice samples, which were specifically extracted from the interior of the megacryometeor fragments. Finally, in order to determine the biological contamination of the ice, three aliquots of ice-melt water were sampled from the surface of the megacryometeor and plated onto solid media with nutrient agar (Panreac Cultimed) for selective growth culture. Subsamples were removed for PCR amplification of 16S rDNA. The subsequent cloning and sequencing of the samples, using a 3130 Genetic Analyzer and the Microseq software (Applied Byosistem), revealed the following species: Brevundimonas intermedia, Kocuria rosea, Achromobacter sp., Sphingomonas sp., Pantotea sp. and Acinetobacter sp. Basically, the circumstances surrounding the Mejorada del Campo incident and the hydrochemical and isotopic features determined in the ice samples are in a good agreement with previous results6,10,11 concerning other megacryometeors. It is well known that precautionary principles require that environmental managers should be prudent when making decisions, where there is still an incipient understanding about the underlying scientific issues. As previously defined, although several hypotheses have been advanced, no geophysical model is able to adequately give explanation of what factors cause the ice This journal is ª The Royal Society of Chemistry 2008

nucleation and growth, or how megacryometeors can be actually formed and maintained in the atmosphere.6,9–11 However, it is a fact that tropospheric ice chunks, weighing tens of kilograms, do fall provoking verifiable hazards without a clear knowledge about the environmental implication of the whole process that rules their formation in the atmosphere (more or less anthropogenically related). This article offers new data about the Mejorada del Campo megacryometeor and documents the value of having both civil and scientific institutions involved in conducting follow-up investigation of such atmospheric ice falls. Institutions should be ready and alert to the need for proper environmental and logistic responses to these incidents.

Acknowledgements The authors acknowledge the collaboration of the workers and director of Iberostil S.L., and the institutional support provided by the Criminalistic Service of the Guardia Civil and CSIC. Also, thanks to Dr David Hochberg for the revision and correction of the English version. Two anonymous referees are acknowledged for their helpful comments that improved the original manuscript.

References 1 National Environmental Monitoring Initiative, US Environmental Protection Agency, Integrating the Nation’s Environmental Monitoring and Related Research Networks and Programs, 2006, http://www.epa.gov/cludygxb/Pubs/factsheet.html. 2 http://www.igac.noaa.gov/meetings.php. 3 J. Martinez-Frias and J. A. Rodriguez-Losada, in Comet and Asteroid Impacts and Human Society, International Council for Science (ICSU), ed. P. Brobovsky and H. Rickman, Springer, 2007, pp. 339–352. 4 International Union of Geological Science, Geoindicators, http:// www.lgt.lt/geoin/. 5 J. Martinez-Frias, F. Lo´pez-Vera, N. Garcıá, A. Delgado, R. Garcıá and P. Montero, Geotimes, News Notes, American Geological Institute, Atmospheric Sciences, June/2000, http://www.agiweb.org/ geotimes/june00/hailstones.html. 6 J. Martıńez-Frıás, M. Millań, N. Garcıá, F. Lo´pez-Vera, A Delgado, R. Garcıá, J. A. Rodrı´guez-Losada, E. Reyes, J. A. Martıń Rubı´ and A. Go´mez-Coedo, Ambio, 2001, 30-7, 450–453. 7 X. Bosch, Science, 2002, 297, 765. 8 J. Martinez-Frias and D. Travis, in Environmental Catastrophes and Recovery in the Holocene, 2002, Brunel University, Uxbridge, UK. 9 K. Brink, D. Travis and J. Martinez-Frias, Wisconsin Geographical Society, 57th Annual Meeting, September 19–20, 2003, UW-Eau Claire, USA. 10 E. Santoyo, R. Garcıá, J. Martıńez-Frıás, F. Lo´pez-Vera and S. P. Verma, J. Chromatogr., A, 2002, 956, 279–286. 11 J. Martinez-Frias, A. Delgado, M. Millan, E. Reyes, F. Rull, D. Travis, R. Garcıá, F. Lo´pez-Vera, J. A. Rodriguez-Losada, J. A. Martıń-Rubi, J. Raya and E. Santoyo, J. Atmos. Chem., 2005, 52, 185–202. 12 J. Martinez-Frias and A. Delgado, Ambio, 35-6, 314–316. 13 M. M. Joshi, A. J. Charlton and A. A. Scaife, Geophys. Res. Lett., 2006, 33, L09806, DOI: 10.1029/2006GL025983. 14 D. T. Shindell, Geophys. Res. Lett., 2001, 28(8), 1551–1554. 15 S. J. Oltmans and D. J. Hofmann, Nature, 1995, 374, 146–149. 16 Megacryometeors. http://tierra.rediris.es/megacryometeors. 17 H. Craig, Science, 1961, 133, 1702–1703. 18 J. Saul, Int. Q. Meteorit. Meteorit. Sci., 2007, 12(2), 20–21. 19 M. Beech, Int. Q. Meteorit. Meteorit. Sci., 2007, 12(4), 17–19. 20 P. Spichtinger, K. Gierens, U. Leiterer and H. Dier, Meteorol. Z., 2003, 12-3, 143–156. 21 P. A. Hirschberg and J. M. Fritsch, Mon. Weather Rev., 1991, 119, 518–550. 22 G. T. Meaden, J. Meteorol., 1977, 2, 137–141.

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This journal is ª The Royal Society of Chemistry 2008

PRESERVATION WINDOWS FOR PALEOBIOLOGICAL TRACERS IN THE MARS GEOLOGICAL RECORD DAVID C. FERNÁNDEZ-REMOLAR1, OLGA PRIETOBALLESTEROS2, CÉSAR MENOR SALVÁN1, MARTA RUÍZ BERMEJO1, FELIPE GÓMEZ1, DAVID GÓMEZORTIZ2 AND RICARDO AMILS1,3 1 Centro de Astrobiología, INTA-CSIC, Ctra Ajalvir km. Torrejón de Ardoz, Spain, 2 3Unidad de Microbiología, Centro de Biología Molecular, Universidad Autónoma de Madrid, Spain, Instituto Geológico y Minero de España, Salamanca, Spain

1. Introduction: a new perspective of the Mars sedimentary record For years, the Mars robotic missions have provided different evidences that Mars had an active hydrologic past which involved the emergence of distinctive sedimentary systems and its corresponding weathering sources. Minor geomorphic features to regional-scaled geomorphological structures have been used to infer sedimentary systems as deltaic, fluvial, lacustrine or marine-like environments (REFs Malin and Edget) to have occurred sometimes in the Mars history (Carr, 2006). In this context, the information obtained by geomorphological interpretations have inferred those physical conditions –e.g. hydrological activity, water energy or climatic evolutionthat were in equilibrium with the landforms (Baker, 2001). In recent times, new instrumentation aboard the planetary missions to Mars (e.g. IR specs in the Mars Oddyssey, Mars Express and MRO), as well as APXR and Mössbauer specs of MERs have shed light in the mineralogical and geochemical composition of some ancient materials. Both orbiter and rover explorers have recognized that the two main agedifferentiated Mars terrains have differential mineralogy and geochemistry (Poulet et al., 2005; Bibring et a., 2006). In the oldest Noachian areas (age older than 3.8 Ga), the Mars Express orbiter has detected phyllosilicates concentrated in layered terrains (Michalski and Noe Dobrea, 2007), currently covered by younger deposits of lavas and other sediments. On the other side, Late Noachian to Herperian younger terrains (3.8-3.0 Ga) are composed of sulfates, some of them bearing iron (Squyres et al., 2004; Morris et al., 2004; Fernández-Remolar et al., 2005). Different mineralogies have been used as indirect evidences in inferring the hydrogeochemical processes occurring on the Mars surface, which are the result from the interaction between climate, hydrosphere and geosphere ogf

Mars. On earth, phyllosilicates are formed in diverse environments including hydrothermal, sea and lacustrine through precipitation, sedimentary through weathering and later transportation or metamorphic transformation through strong changes in temperature and pressure. Thus, the clays found on Mars, if are sedimentary (Michalski and Noe Dobrea, 2007), denote high rates of weathering driven by CO2-saturated meteoric waters (Francois and Walker, 1992; Franck et al., 1999) under warm conditions, which induce aqueous acidification through CO2 hydration to H2CO3 (Orr et al., 2005). On the other hand, the hydrated sulfates bearing ferric iron precipitate under strong acidic brines (pH < 3), which are sourced in the sulfide weathering by oxygen-rich meteoric waters and/or anaerobic iron oxidizers (Amils et al., 2007; Fernández-Remolar et al., 2008c). As a result, an association of different iron-bearing sulfates as copiapite, jarosite and schwertmannite co-occurs with any other sulfates like gypsum or epsomite formed by cations sourced in the silicate dissolution. Later on, the occurrence of Hesperian to Amazonian outlow channels and great catastrophic landforms suggest that Mars had some postNoachian planetary events of thermal reactivation and transient water masses (Carr, 2006) when great iced terrains –permafrost and glacier-like systems- were melted. Such episodic massive release of water took place under a declining atmosphere that was transitionally recovered through the volatile replenishment (Baker, 2001). As a consequence, different highland lacustrine and lowland sea-like regions were created or reactivated during one or several episodes where the climatic conditions were warmed up. If life arised sometimes in early Mars, it should have adapted and evolved to the long-term climate evolution that is driven by the inner planetary activity. Moreover, preservation of biological information produced by the Mars life into the Mars geological record must have occurred according to fossilization processes that depend on the crust diagenetic geochemistry, which also emerges from atmosphere, crust composition, hydrological activity and heatflow. According to the Mars geological record uncovered along different planetary missions, we propose some preservation windows in which primary biological information may have been recorded in form of one or several fossil states ranging from pure organic compouds to resistant mineralized remains. The following preservation windows are considered to have preserved paleobiological entities:     

Early Noachian phyllosillicate deposits Fluvial to lacustrine or marine-like Early Noachian materials Hydrothermal-associated deposits Late Noachian to Hesperian sulfate to hematitic basaltic sands. Post-Noachian fine-grained lacustrine or marine to fluvial-cold climate deposits

To these five preservation windows two more can be added in relation to subsurface regions and rock coatings. The main reason is whatever the surface conditions have reigned on the Mars surface, its subsurface counterpart may have had other more stable concerning to temperature and shielding against radiation. Moreover, environmental coditions are easier to control by microbes in the susbsurface as it has been showed concerning to temperature and pH (Gómez et al., 2004; Fernández-Remolar et al., 2008b). Under these favorable circunstances, biogeochemical cycling can operate during the long and cold episodes that followed the benign climate inferred for the Early Noachian age. Subsurface areas in close approach to volcanic centers with hydrothermal activity are exposed to high mineralization rates that is an essential parameter for morphological conservation. Finally, given that Mars has developed along its long history fluvial and desertic systems in which boulders of sedimentary bars or pavemented soils are covered by complex thin films composed by oxides, sulfates, carbonates and weathering silicates (Potter and Rossman, 1977; Giorgetti and Baroni, 2007). Microbial endoliths and endolithic structures (Golubic and Schneider, 2003), not discussed in this work, should be added to these geobiogical entities of great importance for searching Life on Mars.

2. Taphonomic and organic chemistry-related processes involved in preservation Paleobiological remains and fossils are currently concerned as real biological entities, but are not life entities. On the contrary, they result from the interplay and succession of several geo-biological processes that occur before, during and after it is buried, and which are recorded additionally to the primary remain (Fernández-López, 1991, 1995; Brocks and Summons, 2005). As a consequence, fossil entities record, not only some information concerning to its biological origin, but also all those processes that have played any role in generating the preserved entities. A good example is the organics obtained in sedimentary rocks that come from the multi-way degradation of primary biomolecules under different thermal and compositional regimens along the rock diagenesis (Brocks and Summons, 2005). In this sense, exposition of biomolecules to iron- and sulfur-rich environments has a strong imprint in the final geopolymers that are associated to iron and sulfur (Sinninghe Damsté and de Leeuw, 1990). Therefore, the fossilization process can follow complex pathways that involve preservation before and after definitive burial, fossilization phases known and biostratonomy and fossildiagenesis, respectively (FernándezLópez, 1991, 1995). Obviously, the final fossil state will be the result of all these stages and which can be so simple than a fast and in-situ burial that is the best case for the preservation of chemical fossils. The parameters involved in the formation of the preserved entities (fossils and/or any paleobiological trace) are countless (Farmer and Des

Marais, 1999). They range from molecular processes (Banfield et al., 2005) currently driven by microbes to planetary-scaled events as sea level global changes (FernĂĄndez-LĂłpez, 2007). These planetary events are ruled by macrotectonic to long-term climatic changes affecting global biogeochemical cycles (Brasier, 1992) that are drivers of preservation in many cases like redox potential to organics preservation. Whatever the processes drive the fossil preservation, all micro and macro mechanisms are sustained on some few physicochemical parameters which are essential in the final record of the biological information. Essentially, these parameters are hydrodynamics (diffusion vs. advection), temperature (biogeochemical reaction, mineralization and recrystallization rates), redox potential (oxidation processes), and ion concentration (mineralization), which dominance or co-occurrence can diversify the paleobiological record into different preservation windows. Obviously, exceptional fossil preservation -e.g. the so called conservation deposit fossil-lagerstĂ¤tten as Burguess Shale (Conway Morris, 1990; Seilacher, 1990)- result from the positive concurrence of all these parameters; but from the interplay of all parameters will emerge the diversity of preservation windows that enrich the geological record on Earth. In the next section some ancient and modern Terrestrial analogs will be considered to exemplify different kinds of preservation windows that can be expected to occur in the extensive geological record of Mars. Theoretical characterization of preservation windows based on solely ion the four parameter interplay can be previously carried on as showed in Figure 1. Oxygen availability and redox potential are two elements that rule the molecular preservation of life. However, it should be noted that low oxygen fugacity does not mean low redox potential, which can be essential to understand preservation pathways on Mars. Although biotic and abiotic oxidation destroys most of low resistant biomolecules (e.g. sugars, proteins and nucleic acids), a minor part transforms to macromolecular humic complexes and geopolymers precursing kerogen and bitumen. On the other hand, more resistant fatty acids and lipids are transformed to geolipids and hydrolyzed hydrocarbons, but maintaining the original structure that enable an easier identification concerning to its biological origin (Brocks and Summons, 2005). In any case, it has to be noted that not only redox and other primary factors but temperature during diagenesis is an essential factor to maintain preserved the molecular traces of life. Indeed, organic matter suffers thermal destruction under metamorphism, show a distinctive preservation under high-temperature extreme areas that are close to hydrothermal centers (Brocks and Summons, 2005) as is observed in the Upper Devonian Tharsis deposits shortly in the next section. On the contrary, when the thermal and redox history of the preserving remains converge into a positive way, exceptional conservation of biopolymers occur and some of them can be amplified using current molecular biology techniques (Logan et al., 1993).

A final but not less important factor to consider as essential to understand the processes involved in preserving biology is time. Indeed, rock aging, currently known as diagenesis, encloses strong changes in several parameters that are drivers in selecting some paleobiological entities. In this sense, oxidant replenishment in sediment fluids produces changes in the hydrochemistry exposed to meteoric waters that produces the complete oxidation of the organic matter (FernĂĄndez-Remolar and Knoll, 2008) to heavy carbon-bearing compounds of complicate determination.

Figure 1. Theoretical representation of preservation windows (A to D) displayed in a threedimensional space defined by essential environmental parameters for preservation of biological information as temperature, eH and ion concentration (named as S in mgÂˇL-1). (A) represents ion-enriched medium temperature and high eH conditions compatible for acidic to neutral environments where paleobiological structures are preserved under a fast mineralization. Same S and eH records changing to high temperature conditions (B) would be a window for hydrothermal areas where high mineralization rates are also accounted. (C) corresponds to medium-temperature and low-concentrated fluids under reducing conditions that favor a net preservation of organics. Note that some taphonomic gradients emerge according to changes in environmental parameters as observed in RĂo Tinto for eH, pH and ion concentration (see Fig. 2).

3. Preservation in terrestrial analogs Some Terrestrial habitats dated from Archean to modern ages have been claimed as feasible analogs of different Mars potential habitats that have emerged over time. Such a statement is based on a methodological background that considers the Terrestrial life inhabiting Earth regions as

the unique reference to detect life in other regions of the Universe. Moreover, given that there is not a conceptual base to define Life, the Terrestrial nature has currently provided the source to infer what particular Life inhabites a defined area considered as Mars analog. On the other hand, as one of the main connections between life and habitable region are water, it can be deduced that water is the main fact that characterizes a Mars potential habitat. As a result, water, as the exchange matrix for matter and energy used by life, and life itself are intrinsically linked in a search for extinct or extanct non-terrestrial living beings on Mars. Remaining the water as the general fact basing a potential habitat, many other enviromental parameters as pysichochemical or hydrogeochemical produce differenciation between environments, and which will also cause its characteristic imprint in preserving paleobiological traces. Despite of some Terrestrial analogs are inferred through the geological record, they bear information to unlock some essential questions to uncover the paleoenvironmental conditions that could have been linked to the emergence of life on Mars. The main reason is due to the atmospheric system of both Mars and Earth planetary bodies can have same volatile inventory, which interacted a low-differentiated crust but evolving lately to different global geochemistries (Fernández-Remolar et al., 2008a). Therefore, very early geohistorical Terrestrial analogs as Isua metasedimentary sequences older than 3.7 billions of years (Rosing et al., 1996; Fedo, 2000) are of great importance to trace back those driving forces that might have impeled the possible emergence of life on Mars. Later on, the geological record of subsequent Archean environments (Pilbara and Barberton Archean deposits younger than 3.7 billion of years) are reference systems to determine those mechanisms driving the planetary divergence between Mars and Earth (Grady and Wright, 2006). Table 1 displays an intend of equivalence between some Mars potential habitats and its Terrestrial analogs that mention the driving environmental parameters involved in preservation. As showed in Table 1, combined study on ancient and modern environments will improve in acquiring knowledge not only about that processes involved in recording paleobiological information, but also those involved in long-term preservation. Therefore, for unlocking the aging processes that favor preservation of some paleobiological over other entities it is necessary to appeal to analogous modern environment. In few cases the geological record of ancient and its corresponding modern equivalent coexist within the same area as seen in the Río Tinto fluvial basin (Fernández-Remolar et al., 2005; Fernández-Remolar and Knoll, 2008). All environments described in Table 1, as other Terrestrial not mentioned here, deserve a detailed analysis in order to have a perspective for the preservation evolutionary changes occurred on Mars since its earliest evolutionary stages. This would demand indeed an extensive work dealing on characterization of the taphonomic processes driving in the

preservation of any paleobiological entitie that can be potentially produced in each analog. On the contrary, some interesting analogs having a Mars real couterpart are herein described in order to provide an idea how the research on the preservation in Earth analogs is essential to drive the exploration for extinct Life on Mars.

TABLE 1. Correspondence between Mars Potential Habitats and its Terrestrial counterparts showing some parameters that can favor preservation. Note that the Isua metamorphic materials are topic of a intense debate concerning whether â&#x20AC;&#x153;the polimictic meta-conglomeratic unitâ&#x20AC;? (Fedo, 2000) is in fact a true sedimentary deposit.

3.1 PRESERVATION DEPOSITS

FLUVIAL

MARINE

ARCHEAN

Early terrestrial habitats are likely the most suitable approaches to the early Mars environments (Fig. 2A-B). During the Archean, the upper part of the Terrestrial crust, uncovered during billion of years by terrestrial flora, was direcly exposed to a CO2-driven atmospheric weathering on basaltic-like primary rocks that induced high physicochemical alteration. As a result, detrital sediments were massively transported to the marine environments through transition sedimentary systems like deltaic. Moreover, the aggressive chemical weathering introduced silica into solutions (Hamade et al., 2003) in major proportion than the subaqueous geothermal systems. On the other hand, some oxidative processes might have occurred (Ohmoto, 2004) impeled by photochemical pathways affecting to ferrous iron in form of aqueous complexes that were sourced in hydrothermalism. Such oxidation likely resulted in the production of ferric deposits as Banded Iron formations and red beds. Interestingly, carbonates did not occurred before 3.5 billion of years (Grotzinger, 1994) despite of the early Earth atmosphere was mainly composed by CO2 that suggest some acidification mechanisms preventing the massive production of them. Strong redox gradients, as well as fast silicification and ferruginization are mechanisms expected to favor preservation of organics and paleobiological morphologies, respectively (Walsh and Westall, 2003). An exceptional case to be mentioned is the preserved microbial mats occurring in of deltaic tidal flat deposits of the 3.2 billion-of-year Mesoarchean Moodies Group in Barberton (Noffke et al., 2006), South Africa (Fig. 2 CD). Higher photochemical activity producing oxidants, but under an atmosphere lacking oxygen, as well as high sedimentation rates might have favored strong redox gradient driving to reducing conditions in shallow areas of the water column. Moreover, abundance of silica in sediment porewater likely increased the preservation potential of microbial remains due to a fast mineralization of the deposits. Although organics would have prevailed in these conditions, late metamorphic processes related to a combining increase of sediment pressure and temperature (Brocks and Summons, 2005) have favored the organic destruction. Analogous Mars environments can be found in shallow deltaic deposits that occur elsewhere infilling crater basins and different-sized impact craters of Noachian age (Fasset and Head, 2005). 3.2 PRESERVATION IN HYDROTHERMAL ACTIVITY

ENVIRONMENTS

LINKED

Hydrothermal activity has been present to Mars since early Noachian to recent times. Such a process has been recognized through the volcanism (Fig. 3A) that affected to the crust water mobilization in form of lacustrine and fluvial systems (Cabrol and Grin, 2001; Schulze-Makuch et al., 2007),

as well as silica-rich deposits discovered recently in the Gusev Crater (Prof. Raymond Arvidson personal communication showed in Fig. 3B), which likely originated under acidic hot fluids.

Figure 2. Noachian deltaic deposits (Grant et al., 2007) infilling the Holden Crater in Mars (aC) and preserved microbial mat structures in Archean deltaic materials (D-E) at Barberton Greenstone Belt, South Africa. (A) Mars Orbiter Camara (MOC) wide angle image of the Holden Crater area where a deltaic structure occurs (white rectangle). (B) THEMIS Visible Image V17376003 (Themis Public Data Releases, Planetary Data System node, Arizona Srate University at http://themis-data.asu.edu) showing the deltaic structure filling the crater. (C) High Resolution Imaging Science Experimet (HiRISE) image PSP_001468_1535 (supported by the NASA/JPL/ASU at http://hirise.lpl.arizona.edu) boarded at The Mars Reconaissance Orbiter (MRO) unlocking the deltaic-like sequence that were sedimented inside the crater.

On Earth, geothermal systems areas into volcanic centers (Fig. 3C), but also to hot springs that do not manifest as volcanic buildings, and much difficult to detect using remote sensing, are other environments, are widely spreaded. Currently, they favor high silica and sulfur fluxes that induce high mineralization rates of a very resistant mineral-complexes as opaline silica currently forming sinters (Schinteie et al., 2007). Such a fast mineralization induces a instantaneus silicification of living microbes (Jones et al., 2001) that facilitates preservation of organics over time. A rapid organic enclosing preventing its subsequent degradation occurring under exposure to oxidative surface conditions. In some cases, silicification in geothermal systems produces exceptional replica of the microbial components(Fig. 3D) taking part in the community. Interestingly, silica preservation in some geothermal systems as the Parakiri Stream sourced in the Rotokawa Geothermal of New Zealand (Schinteie et al., 2007) occur under acidic condition, which will be described in the following section. On the other hand, hot spring carbonate deposits have been also described to preserve biological information (Kazue, 1999). Although carbonates are pervasive on the surface regions of Mars, geothermal systems driven by CO2-rich solutions could be present in some areas of the planet. 3.3 PRESERVATION IN ACIDIC ENVIRONMENTS Late Noachian to Early Hesperian surface environments on Mars were probably ruled by global acidic conditions as inferred through the sulfate and oxide materials occurring in different areas as Meridiani (Fig. 4A), Gusev Valles Marineri and North Polar regions. However, acidic environments have been longly considered incompatible to live and related to highly contaminated areas (Blowes et al., 2005) given that some acidic solutions are sorced in mine tailings resulting from mining operations (Davis et al., 2000; Blowes et al., 2005). However, recent studies on the Earth geological record of ancient and modern acidic environments have spotted several systems to be natural (Fernández-Remolar et al., 2005; Benison, 2006). Preservation in modern and ancient acidic environments have been reported from the Tyrrel Lake at nothwestern Victoria and other lacustrine areas of Western Australia (Benison and Laclair, 2003), acidic mine drainage of Indiana (Brake et al., 2002) and Río Tinto (Fernández-Remolar et al., 2007; Fernández-Remolar and Knoll, 2008). Combination of studies on modern and ancient environments are essential to understand long-term preservation in acidic environments. Interestingly, in Rio Tinto modern and ancient sediments dateding back to more than 6 million of years coexist in the same area (Moreno et al., 2003; Fernández-Remolar et a., 2005), which favor an intregrative study concerning to the syngenetical and postsedimentary taphonomic processes over time (Fig. 4C-D). Moreover, Río Tinto basement drillling has allowed to sample the aquifers sourcing and storing the surface acidic solutions (Fernández-Remolar et al., 2008b).

Figure 3. Context image of the Apollinaris Patera shield volcano (A) showing geomorphic structures related to cladera collapse, sediment infilling and water activity as fan-channeled pattern emerging southwards out from the crater, as well as wall a strong volcano erosion operated by gullies. Such a volcanic building was likely the scenario of hot fluid production displayed as geisers, volcanic chimneys or any other geothermal phenomenon (image was composed using the PIGWAD GIS Mapping system, courtesy USGS Astrogeological Research Program at http://astrogeology.usgs.gov). (B) image PIA09491 silica-rich soil uncovered by Spirit at Gusev Crater that might be interpreted as a consequence of geothermal mineralization of silica-enriched fluids or a strong leaching on basaltic precursors inducing a secondary silica enrichment (courtesy of NASA/JPL-Caltech). (C) Crater caldera of the Kilauea volcano (Island of Hawaii) showing geothermal activity that produce sulfur and silica rich deposits around (withe arrows) (Credited by Prof. Raymond Arvidson and Thomas Stein). (D) Silicified bacili-like microbial structures in sinter deposits of the Rotokawa Geothermal Field in New Zealand (Credit: Dr Richard Schinteie).

Physichochemical and geobiological analysis of the subsurface areas have depicted a very different environmental conditions that change to neutral and reducing with a lower concentration in ions (FernĂĄndezRemolar et al., 2008a). Under these new environmental circunstances geopolymers and specifically geolipids of different origin (Fig. 4D) are preserved. On the contrary, the organics detected in the RĂo Tinto acidic sulfate deposits should have a long-term preservation if protected against the meteoric and diagenetic solutions as demonstrated by Aubrey et al. (2006), which report organics from jarosite mineralogies in the Panoche Valley (California)with an age of 40 million of years. Preservation of sulfates on some Mars deposits (Fig. 4A) like jarosite (a very soluble mineral phase) suggests that late meteoric and diagenetical solutions did only remobilized partially the sulfate and iron to form hematite concretions. Translating this to the Burns Formation (Fig. 4A), assuming that life emerged sometimes on Mars, it could be expected to find organic and paleobiological structures in the same materials. 3.4 PRESERVATION IN THIN FILMS COVERING ROCKS Water history of Mars suggests that rock coatings, desert varnishes and weathering rinds should be present in any sedimentary deposit or embedded inside alteration materials derived from ancient and recent aqueous activity. Spirit MER have endeed provided direct evidences of rock coatings of unequivocal aqueous alteration (Haskin et al., 2005). In this sense, location and analysis of rock coatings of several ages can be essential to trace back the climatic and environmental history (Don and Dickinson, 1989; Liu and Broecker, 2000) of the planet since its origin. In fact, lamina accrection and substrate weathering currently work under seasonal climatic regimen that result from the periodic water availability and thermal conditions. Such an information can obvioulsy essential to determine the ancient to recent surface and subsurface environmental water patterns that took place in Mars. In desert varnishes thin aqueous films contacting the rock surfaces are currently oversaturated in ionic species as silica, manganese or iron remobilized from the rocky substrate (Perry et al., 2006), which favors microbial activity (Kuhlman et al., 2006) and later preservation to organic traces when the conditions are favorable (Perry et al., 2006). In rock coatings and weathering rinds oversaturation under aggressive acidic or alkaline conditions are also favorable conditions to induce high mineralization rates to preserve from microbial strutctures to organics. As a result, a complex mixing of different mineralogies as iron and manganese oxides, sulfates, carbonates, opaline silica and different phyllosilicates (Potter and Rossman, 1977) occur as laminae enveloping weathered rocks. Several environmental conditions beside of desertic can be imprinted in the surface rinds. Some volcanic emissions centered in geothermal activity produce SO2-rich adicic fog which acting on volcanic tephra induce the

formation of silica and sulfate laminae (Schiffman et al., 2006). Ancient fluvial deposits are currently composed of conglomeratic materials which pebbles may show lamina covering depending on the climatic conditions that originated them. The Triassic Buntsandstein fluvial deposits in association to contemporaneous and older lacustrine and aeolian desert-like sedimentary materials, contain rounded pebbles that are covered by ironrich coatings recording paleoclimatic information of great interest (Fig. 5A). Concretely, the RĂo Tinto geological record dating back from Tertiary also shows iron rich coatings which origin is undoubtly associated to seasonal activity of acidic environments (Fig. 5B), and with clear traces of biological activity (Fig. 5C-D). It should be overemphasized the importance of these microdeposits as recorders of modern and ancient biological activity, which can be easily detected in many planetary regions of Mars. As showed by Kuhlman et al. (2006), rock varnishes, as many other film coating environments, are microhabitats inhabited by diverse microbial communities having up to 108 microorganisms per gram of varnish lamina. Such a microbial activity can be traced through biochemical and organic compounds (Perry et al., 2006) that are the base for the development of an exploration strategy for searching life on Mars.

4. Conclusions: Mars preservation windows and strategy for planetary exploration Integrative research on preservation of biological information in terrestrial analogs is essential to build up a consistent strategy to search for extinct life on Mars. From the scientific point of view, any exploration strategy developed for this compelling objective has to deal the diverse Mars geological record which shows different preservation potential of biology depending on the paleoenvironment and diagenetical processes that have conformed it. As a result, different paleobiological entities may have potentially persisted and which detection demand distinctive explorative procedures, sampling techniques and instrumentation (Farmer and Des Marais, 1999). Therefore, the application of preservation windows concept can be of great utility to define a specific explorative strategy based on the preservation potential of a given geological unit. The RĂo Tinto Mars analog can be claimed to illustrate shortly this assumption. As mentioned above, whereas the Rio Tinto surface environment favors preservation of morphologies and organics in iron- and sulfate-rich materials, organics are only preserved in the reducing subsurface; however, preservation is reset to simple preservation of morphologies when all these materials are exposed to a 2-million-of-year diagenesis. Under these varying conditions, the detection of extinct life on Mars would require different instrumentation

going from optical to analytical instrumentation which application strongly drives the exploration strategy of a given area.

Figure 4. Burns Formation at the Burns Cliff in the Endurance Crater of Meridiani (A), Pan cam image PIA03241 resulted of a false color composite mosaic (courtesy of NASA/JPLCaltech) which Whattanga and Wellington sedimentary boundaries (Grotzinger et al., 2005), showed as 1 and 2, respectively; grained-sulfates and the hematites in concretional structures (blueberries) have been identified as mineralogies giving some clues concerning to the acidic conditions of Meridiani Mars area explored by the Opportunity (B) microbial remains embedded in a cryptocristalline matrix composed of SO 4 and Fe3+-bearing mineralogies. (C) 2.1 million of year goetithe layer showing preserved filaments inside having cryptocristalline habit. (D) Organic association preserved in neutral and reducing underground areas of the RĂo Tinto system where hopanoids are present.

Figure 5. Iron-rich coatings on boulders (A) embedded in fluvial Iberian Permotriassic deposits (250 million of years), and (B) inside young RĂo Tinto terrace materials (1000 years old). (C) SEM image showing microbial patches covering coated boulders and (D) EDAX microanalysis (spectrum 1 in (C)) showing a carbon and iron enrichment as expected for microbial films associated to watery environments enriched in iron.

5. Acknowledgements We thank to Prof. Raymond Arvidson, Thomas Stein and Richard Schinteie for providing essential information and very illustrative images. Special thanks to the USGS Astrogeological Program, NASA/JPL-Caltech, NASA/JPL/ASU and the THEMIS Public Data Releases which have provided the Mars surface images. This paper was supported by the Project ESP2006-09487 funded by the Department of Science and Education of Spain.

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Con formato: Inglés (Estados Unidos) Con formato: Inglés (Estados Unidos)

FIRST FIND OF KRISTIANSENITE IN SPAIN: COMPARISON WITH THE TYPE SPECIMEN BY NON-DESTRUCTIVE TECHNIQUES P. Prado-Herrero1 J. Garcia-Guinea2 E. Crespo-Feo2 V. Correcher3 C. Menor4 1

MICINN, Albacete 5, Madrid 28027, Spain.(pedro.prado@micinn.es) Museo Nacional de Ciencias Naturales (CSIC), José Gutiérrez Abascal 2, Madrid 28006, Spain 3 CIEMAT.Av. Complutense 22, 28040 Madrid. Spain 4 Centro de Astrobiología (CSIC-INTA)Ctra. Torrejon-Ajalvir, Km 4 Torrejón de Ardoz, Madrid, Spain 2

ABSTRACT We report herein a new find of kristiansenite from a pocket in an intra-granite pegmatite from Cadalso de los Vidrios, near Madrid, Spain. This specimen of a late hydrothermal scandium silicate has been studied by Environmental Scanning Electron Microscopy with Energy Dispersive Spectrometry probe (ESEM-EDS), Micro-Raman Spectrometry and ESEM-Cathodoluminescence (ESEM-CL), all of them non-destructive techniques. The sample is a single perfect pyramidal monocrystal found in a small cavity less than one mm across. The experimental chemical, molecular and spectral luminescent information was later compared with the type specimen from Norway and the second find, at Baveno, Italy. Our Raman spectrum matches the spectrum of the Norwegian specimen, with minor variation in the intensity of the peaks; the chemical composition recorded by EDS also shows minor variations. In addition, the CL spectrum displays several narrow peaks, probably associated with REE in Ca positions. The geochemical framework of this new locality, with pegmatite pockets in A-type granites rich in Sc-bearing minerals and other REE, have many similarities with those of Norway and Italy. Keywords: kristiansenite, cathodoluminescence, Raman, REE INTRODUCTION Approved by the IMA in 2000 (#2000-51), the name kristiansenite honours Mr. Roy Kristiansen (born in 1943), a Norwegian mineralogist who first noticed the new mineral collected from the Heftetjern pegmatite (of amazonite“cleavelandite” type) 200–250 m long and 3–4 m wide, in the Tørdal area, Drangedal, Telemark County (southern Norway), the

Estudos Geológicos v. 19 (2), 2009

type locality for this disilicate species. It is a triclinic C1 (pseudomonoclinic) sorosilicate, with theoretical formula Ca2ScSn(Si2O7)(Si2O6OH), described as new mineral species by Raade et al. (2002); the crystal structure was determined by Ferraris et al. (2001). The second ﬁnd of kristiansenite was described in the Miniera Seula (Montecatini quarry), Baveno, Verbano–Cusio–Ossola, Piemonte, Italia (Guastoni & Pezzotta,

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2004). The third find, described here, is the specimen found in the granite quarries of Cadalso de los Vidrios, near Madrid, Spain. Scandium ore deposits are uncommon, and mineral species containing Sc as important constituent are scarce. Scandium is more abundant in the Earth’s crust (21,9 ppm, Rudnick and Gao, 2003) than other common elements such as tin, ten times less abundant in the lithosphere. Scandium is widely dispersed in rock-forming and accessory minerals, being mainly associated with ferromagnesian minerals such as pyroxenes and amphiboles. For this reason, a small fraction of Sc is mobilized in pegmatite-related fluids; Sc is present in minerals such as helvite, milarite, epidote, and garnet. It is uncommon to find minerals in which Sc is essential constituent; in fact only nine minerals have been reported. Raade et al. (2002) explained that the incorporation of Sc in minerals is based on the similarities of the outer electronic structure of Sc, Y and some REE, in addition reflects the relatively small size of Sc3+, compared with Y3+ and Yb3+. This characteristic leads to its substitution for trivalent ions, coupled substitutions, or its incorporation in Nb and Ta minerals. GRANITE OF CADALSO DE LOS VIDRIOS Open working quarries produce granite blocks for ornamental purposes in Cadalso de los Vidrios, in northern Madrid, Spain. These granites were uplifted during the Hercynian Orogeny, 380–320 M.y. ago, as well as the associated metamorphic rocks (migmatite, gneiss, schist). Subsequently, they were exposed and partially eroded during the Cretaceous (120–70 M.y.) transgression, and later on, during the Alpine Orogeny

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(ca. 25 M.y. ago), they rose, developing a NE–SW-trending mountain ridge known as the Spanish Sistema Central. The granites of this area are leucogranites, fine to medium grained, with an aplitic texture combined with granophyre, without phenocrysts (Pérez-Regodon, 1970). The rocks consist of orthoclase, quartz, biotite and plagioclase, as the most important mineral constituents. Pegmatite veins and cavities are rarely found in these granites, as otherwise, they would not be used for ornamental uses. They are centimetric veins, in some cases exhibiting elongate cavities with associations of unusual minerals, such as kristiansenite. The periodic inspection of the quarries of Cadalso de los Vidrios, which cover around 7 km2, has yielded several samples from these pegmatites and cavities, with a wide parageneses of minerals: quartz (hyaline, smoky and amethyst), opal (hyalite), microcline and orthoclase, albite (including the cleavelandite habit), muscovite, biotite, chlorite, topaz, tourmaline (black and blue), garnet (spessartine), helvite, zoisite–clinozoisite, bavenite, beryl, fluorite, sphalerite, pyrite, pyrrhotite, arsenopyrite. bismuthinite, chalcopyrite, molybdenite, ferberite, hematite, apatite, fayalite, pyrolusite, scheelite, axinite, titanite, laumontite, prehnite, stilbite, apophyllite, laumontite, chabazite, milarite, aurichalcite, malachite, uranotile, metatorbernite, torbernite, calcite, aragonite, pyrolusite, gadolinite(Y), allanite-(Ce), kamphaugite-(Y), britholite-(Ce), and kristiansenite, the focus of this article. EXPERIMENTAL PROCEDURE The kristiansenite specimen was firstly analyzed in the Inspect-S ESEM of the FEI Company of the Museo

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P. Prado-Herrero et al.

Nacional Ciencias Naturales in Madrid. The chemical analyses were performed with an EDS probe at 30 kV and focal distance 10,4 mm on a large sample, ca. 5 × 5 cm2 without coating. CL spectra and monochromatic CL images were performed in the ESEM with a direct optical coupling to a chamber-mounted Gatan MonoCL3 monochromator. The excitation for CL measurements was provided by a 30 kV electron beam. The emission current ranges between 52 and 68 mA, and the photomultiplier voltage of the PA3 ampliﬁer was 1000 volts for the luminescence spectrum. The CL emission of kristiansenite was analyzed in lowvacuum mode without coating. The micro-Raman spectra were acquired with a ThermoScientiﬁc DRX Raman microscope, which has a point-and-shoot Raman capability of one micrometer spatial resolution. We used a green 532

nm laser with a 100% power of 10 mW, exposition time to laser 60 s, and two acquisitions each. The used objective was 50×. RESULTS AND DISCUSSION The specimen found in a cavity in the Cadalso de los Vidrios pegmatite is a wellformed crystal, ca. 0.5 mm across, with a vitreous luster; it is translucent, white to slightly yellowish, with the wedge-shaped habit typical for this species and oblique faces. It is a single crystal located in a sample of 5 × 5 square centimeters, with paragenetic quartz, microcline, chlorite, kamphaugite, pyrite and other unknown species containing signiﬁcant amounts of Sc, Y and REE. The physical position of the crystal into the cavity (Fig. 1) makes difﬁcult the non-destructive analytical measurements.

Figure 1. ESEM-CL of the kristiansenite specimen: (a) ESEM image of cavity in which is located the crystal with surrounding chlorite, (b) CL spectrum taken from the top face of the crystal.

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FIRST FIND OF KRISTIANSENITE IN SPAIN: COMPARISON WITH THE TYPE SPECIMEN BY NONDESTRUCTIVE TECHNIQUES

Table 1. Comparison of the EDS analyses of holotype kristiansenite with the Spanish sample.

Oxides

Raade G. et al. (2002) (Mean of 11 analyses)

Na2O

0,41

Raade G. et al. (2004). (Mean of 5 analyses)

Cadalso de los Vidrios (Madrid) 0,66

K2O

0,06

CaO

18,45

19,22

17,31

Al2O3

0,35

0,13

2,7

Sc2O3

8,11

11,74

3,2

Fe2O3

1,98

0,5

4,57

SiO2

40,76

40,05

45,36

TiO2

0,08

0,18

0,93

ZrO2

0,43

0,28

SnO2

27,33

25,42

24,09

0,02

0,64

MnO2 Yb2O3

0,54

Y203

0,01

Nb205

0,29

H2O

2,04

Total

100

99,84

100

Table 1 displays the comparison between our EDS results with the published data of the Norwegian type-locality. One can see that: (i) the SiO2 content is 5% higher in the Spanish case, probably due to the surrounding silicates, (ii) the SnO2 values are very similar in all samples, and (iii) the amount of Sc2O3 is signiﬁcantly lower in the Spanish specimen. The probable replacement of Sc3+ by Al3+ or Fe3+ ions in our specimen could explain the reported differences with respect to the data of Raade et al. (2002). The analyses performed on the associated kamphaugite display an abundance of REE, which could also explain the observed chemical variations as well as the differences between the recorded Raman spectra with respect to the Norwegian material (Fig. 2). The peaks obvious match, but minor variations in intensities and shoulders are observed.

In short, Cadalso de los Vidrios is the third geological area in which kristiansenite is found, after the type locality (Heftetjern pegmatite) and the pegmatite ﬁeld at the Montecatini quarry. The three cases are similar from the point of view of the mineral parageneses, Sc-bearing minerals and other REE. The kristiansenite specimen of Cadalso de los Vidrios, as well as other Sc- bearing minerals present in this pegmatite ﬁeld, commonly include REE in the parageneses and were formed during the late hydrothermal stages of

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the miarolitic pegmatite formation. The mineral parageneses found here exhibits coincidences with the Tørdal case. REFERENCES Ferraris, G., Gula, A., Ivaldi, G., Nespolo, M., Raade, G. 2001. Crystal structure of kristiansenite: a case of class IIB twinning by metric merohedry, Zeitschrift für Kristallographie, 216, 442-448 Raade, G., Ferraris G., Gula, A., Ivaldi, G., Bernhard, F. 2002. Kristiansenite, a

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new calcium-scandium-tin sorosilicate from granite pegmatite in TĂ¸rdal, Telemark, Norway, Mineralogy and Petrology, 75, 89-99 Pezzotta, F., 2004. Kristiansenite a Baveno, secondo ritrovamento mondiale della specie. Riv. Mineral. Ital., 28, 4: 247-251 Raade, G., Bernhard, F. and Ottolini, L,. 2004. Replacement textures involving four scandium silicate minerals in the

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Heftetjern granite pegmatite, Norway. Eur. J. Mineral. 16, 945-950. Garcia, G., Gonzalez, C., Bueno, A. 2004. Cadalso de los Vidrios (Madrid). Bocamina nÂş 14, 12-46. Menor C, Prado-Herrero, P., 2008. Kristiansenita, Britolita-(Ce) y otros minerales raros en el granito de Cadalso de los Vidrios, Madrid. Revista de Minerales 2008/2.

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