- Email: [email protected]
Deep-burial hydrothermal alteration of the Pre-Salt carbonate reservoirs from northern Campos Basin, offshore Brazil: Evidence from petrography, fluid inclusions, Sr, C and O isotopes
Journal Pre-proof Deep-burial hydrothermal alteration of the Pre-Salt carbonate reservoirs from northern Campos Basin, offshore Brazil: Evidence from petrography, fluid inclusions, Sr, C and O isotopes Bruno Eustáquio Moreira Lima, Leonardo Ribeiro Tedeschi, André Luiz Silva Pestilho, Roberto Ventura Santos, Joselito Cabral Vazquez, Jarbas Vicente Poley Guzzo, Luiz Fernando De Ros PII:
S0264-8172(19)30595-1
DOI:
https://doi.org/10.1016/j.marpetgeo.2019.104143
Reference:
JMPG 104143
To appear in:
Marine and Petroleum Geology
Received Date: 3 July 2019 Revised Date:
18 November 2019
Accepted Date: 18 November 2019
Please cite this article as: Lima, Bruno.Eustá.Moreira., Tedeschi, L.R., Pestilho, André.Luiz.Silva., Santos, R.V., Vazquez, J.C., Guzzo, J.V.P., De Ros, L.F., Deep-burial hydrothermal alteration of the Pre-Salt carbonate reservoirs from northern Campos Basin, offshore Brazil: Evidence from petrography, fluid inclusions, Sr, C and O isotopes, Marine and Petroleum Geology (2019), doi: https:// doi.org/10.1016/j.marpetgeo.2019.104143. This is a PDF file of an article that has undergone enhancements after acceptance, such as the addition of a cover page and metadata, and formatting for readability, but it is not yet the definitive version of record. This version will undergo additional copyediting, typesetting and review before it is published in its final form, but we are providing this version to give early visibility of the article. Please note that, during the production process, errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain. © 2019 Published by Elsevier Ltd.
1
1
Deep-burial hydrothermal alteration of the Pre-Salt carbonate
2
reservoirs from northern Campos Basin, offshore Brazil: Evidence
3
from petrography, fluid inclusions, Sr, C and O isotopes
4 a, b
, Leonardo Ribeiro Tedeschi c, André Luiz Silva
5
Bruno Eustáquio Moreira Lima
6
Pestilho c, Roberto Ventura Santos d, Joselito Cabral Vazquez c, Jarbas Vicente Poley
7
Guzzo c, Luiz Fernando De Ros b
8 9
a
Petrobras S.A., Avenida República do Chile, 330, East Tower, Rio de Janeiro, RJ,
10
20031-170, Brazil; [email protected]
11
b
12
Gonçalves, 9500, Porto Alegre, RS, 91509-900, Brazil; [email protected]
13
c
14
Universitária, Rio de Janeiro, RJ, 21941-915, Brazil;
15
[email protected]; [email protected];
16
[email protected]; [email protected]
17
d
18
Central, Brasília, DF, 70910-900, Brazil; [email protected]
19 20 21
Graduate Program in Geosciences, Rio Grande do Sul Federal University, Av. Bento
Petrobras, Research Center (CENPES), Avenida Horácio Macedo, 950, Cidade
Geosciences Institute, University of Brasília, Campus Universitário Darcy Ribeiro, Ala
2 22
ABSTRACT
23
Petrographic, mineralogical, elemental, isotopic and fluid inclusion analyses were
24
integrated to unravel the diagenetic evolution of Brazilian Pre-Salt lacustrine carbonate
25
reservoirs of northern Campos Basin, southeast Brazilian margin. Detailed thin section
26
and cathodoluminescence petrography, scanning electron microscopy and electron
27
microprobe analyses established a paragenetic evolution of diagenetic processes and
28
products, comprising extensive dolomitization, silicification, and dissolution. A
29
paragenesis including saddle dolomite, macrocrystalline calcite, mega-quartz, Sr-barite,
30
celestine, fluorite, dickite, sphalerite, galena, and other metallic sulfides filling fractures
31
and dissolution porosity, and aqueous fluid inclusions with homogenization
32
temperatures of 92 to 152 °C and salinities between 13 to 26 wt. % eq. NaCl
33
characterized a hydrothermal system with some analogy to carbonate-hosted Pb-Zn
34
Mississippi Valley (MVT) and Irish-type deposits. Petroleum inclusions and solid
35
bitumen testify atypical oil generation and migration, associated with the hydrothermal
36
flow. The host Pre-Salt spherulitic and fascicular carbonates present highly radiogenic
37
87
38
Hydrothermal phases show δ18O values more negative than syngenetic and diagenetic
39
carbonates. The δ13C values are interpreted as result of interaction between the
40
hydrothermal fluids and the host rocks. The combined data set provides clear evidence
41
of intense hydrothermal alteration of northern Campos Basin Pre-Salt reservoirs at
42
deep-burial conditions (> 2 km), possibly related to Late Cretaceous or more probably
43
Paleogene magmatic activity. Mixed-sourced fluids bearing a basinal signature fed the
44
hydrothermal system and promoted dissolution of the host rocks. The hydrothermal
45
alterations had strong impact on the porosity, permeability, and heterogeneity,
Sr/86Sr ratios, indicating strong interaction with continental crust materials.
3 46
contributing, together with the associated fracturing, to the excellent production
47
performance of the Pre-Salt reservoirs.
48
Keywords Pre-Salt; lacustrine carbonates; Campos Basin; deep-burial; hydrothermal
49
alteration; magmatic activity
50
1. Introduction
51
Carbonate reservoirs usually have complex and heterogeneous diagenetic
52
evolution as result of their reactive character. Common association with faults and
53
fracture systems often results in intense dolomitization, silicification and/or dissolution
54
(e.g., Gregg and Sibley, 1987; Machel and Mountjoy, 1987; Packard et al., 2001;
55
Corbella et al., 2006; Wilson et al., 2007; Callot et al., 2010; Fontana et al., 2010).
56
Hydrothermal alteration of carbonate rocks has been described in diverse basins, and its
57
characterization is important, both to differentiate from the “conventional” diagenetic
58
evolution of carbonates, and because hydrothermal processes substantially impact,
59
positively or negatively, the quality and performance of carbonate reservoirs (e.g.,
60
Machel and Lonnee, 2002; Davies, 2002, 2004; Machel, 2004; Davies and Smith, 2006;
61
Lonnee and Machel, 2006; Biehl et al., 2016).
62
The hydrothermal
alteration of carbonate deposits involves complex
63
physicochemical processes of interaction with hot fluids, promoting the precipitation of
64
mineral assemblages commonly including saddle dolomite, fluorite, barite, anhydrite,
65
dickite, sphalerite, and pyrite (Neilson and Oxtoby, 2008). These processes indicate the
66
existence of geothermal anomalies, as well as permeable pathways, such as deep fault
67
systems, to allow the percolation and interaction of fluids (Davies and Smith, 2006; Xu
68
et al., 2015; Mansurbeg et al., 2016).
4 69
The Lower Cretaceous Pre-Salt carbonate deposits of the Coqueiros and Macabu
70
formations of northern Campos Basin host important hydrocarbon reservoirs (Dias et al.,
71
1988; Winter et al., 2007). In Pre-Salt deposits there are significant variations in their
72
quality (porosity and permeability) due to the complex diagenesis (Herlinger Jr. et al.,
73
2017; Lima and De Ros, 2019). Recently, evidence on hydrothermal alteration of Pre-
74
Salt deposits has been presented for the Campos Basin (Alvarenga et al., 2016; Vieira
75
de Luca et al., 2017; Lepley et al., 2017; Herlinger Jr. et al., 2017; Lima and De Ros,
76
2019) and for Kwanza Basin, offshore Angola, its western African counterpart (Poros et
77
al., 2017; Girard and San Miguel, 2017; Teboul et al., 2017, 2019). Alvarenga et al.
78
(2016) identified hydrothermal vents in seismic sections of central Campos Basin.
79
Vieira de Luca et al. (2017), Poros et al. (2017), Girard and San Miguel (2017) and
80
Teboul et al. (2019) suggested that hydrothermal fluids ascribed pervasive silicification
81
of Pre-Salt carbonates based on petrographic, isotopic data and fluid inclusions.
82
Herlinger Jr. et al. (2017) and Lima and De Ros (2019) have shown evidence of
83
hydrothermal alteration in Pre-Salt reservoirs from northern Campos Basin, based on
84
the description of hydrothermal mineral phases in thin sections, such as coarse quartz,
85
saddle dolomite, sulfides and sulfates.
86
Although evidence of hydrothermal action in Campos Basin have been discussed
87
by Vieira de Luca et al. (2017), Herlinger Jr. et al. (2017) and Lima and De Ros (2019),
88
a detailed study of hydrothermal mineral phases is needed to improve the
89
characterization of the hydrothermal processes in this and other Pre-Salt areas, and has
90
not been published yet. Particularly, it is important to examine the fracture-related
91
dolomitization, silicification, dissolution, and the precipitation of typical hydrothermal
92
mineral phases, such as saddle dolomite, macrocrystalline calcite, mega-quartz, sulfides
93
and sulfates filling secondary porosity. With this aim, an integrated study was
5 94
performed on four cored wells, comprising detailed petrographic characterization
95
performed through microprobe, electron and optical microscopy, temperature and
96
salinity data from fluid inclusion analyses in diagenetic carbonates, quartz and Sr-barite,
97
and C, O and Sr isotopic analyses in carbonates and Sr-barite. Understanding the
98
processes and patterns of hydrothermal alteration of Pre-Salt deposits is of key
99
importance for reservoir quality assessment during exploration and for the production
100
101
optimization of these extraordinary reservoirs.
2. Geological setting
102
The Campos Basin is situated in southeastern Brazilian margin and covers an
103
area of approximately 100,000 km2, with only 5,800 km2 onshore. The study area is
104
located in the offshore part of the northern Campos Basin (Fig. 1). The stratigraphy of
105
the study area is based on the stratigraphic chart from Winter et al. (2007) (Fig. 2). The
106
reservoirs being addressed here are part of the Coqueiros (rift stage) and Macabu (sag
107
phase) formations from of the Pre-Salt interval of the basin.
108
The rift sequence lies on a basement of the Neoproterozoic Ribeira Fold Belt,
109
which includes a variety of metamorphic and magmatic rocks, such as tonalities,
110
granodiorites, granites and gabbros (Winter et al., 2007; Tupinambá et al., 2012). The
111
initial rift evolution of the basin is represented by Hauterivian volcanics from the
112
Cabiúnas Formation and by Barremian to Lower Aptian sediments of the Itabapoana,
113
Atafona, and Coqueiros formations (basal interval of the Lagoa Feia Group; Fig. 2). The
114
Itabapoana and Atafona formations consist of conglomerates, sandstones, siltstones,
115
arenites of stevensitic ooids and shales (Armelenti et al., 2016). The Coqueiros
116
Formation is composed by bioclastic rudstones and grainstones intercalated with
6 117
dolostones and organic-rich shales (Baumgarten et al., 1988; Dias et al., 1988; Castro,
118
2006; Thompson et al., 2015).
119
The post-rift is defined by the upper interval of the Lagoa Feia Group, during
120
which evolved the so-called sag stage of the basin. In addition to Itabapoana Formation,
121
this sequence comprises the Gargaú, Macabu, and Retiro formations that were deposited
122
during the Middle/Upper Aptian (Fig. 2). The Macabu Formation comprises
123
argillaceous and carbonate laminites as well as spherulitic and fibrous carbonate crusts.
124
These later strata are interpreted as chemical precipitates deposited in alkaline lacustrine
125
environments under arid climate conditions (Wright, 2011, 2012; Tosca and Wright,
126
2014; Wright and Barnett, 2014, 2015; Wright and Tosca, 2016; Herlinger Jr. et al.,
127
2017; Lima and De Ros, 2019).
128
The Retiro Formation (Fig. 2), a thick accumulation of evaporites, was deposited
129
during the late Aptian (Late Alagoas local stage) marine incursion, under arid climate
130
conditions (Leyden et al., 1976; Winter et al., 2007), being composed essentially of
131
anhydrite, halite, and also by bittern salts, such as sylvite and carnallite (Rodriguez et al.,
132
2018). These evaporites provide the stratigraphic seal for the large Pre-Salt hydrocarbon
133
accumulations, as well as for the migration of hydrothermal fluids through deep faults
134
(Lima and De Ros, 2019). Above the evaporites, a carbonate/clastic sequence was
135
deposited from Albian to Recent, which are out of the scope of the present study.
136
Magmatic activity has been recorded in different stages of Campos Basin
137
evolution. Winter et al. (2007) described important extrusive (basalts and hyaloclastites)
138
and intrusive (diabase and gabbro) magmatic events with Ar/Ar ages from 81.5 to 83.2
139
Ma. In the Paleogene, volcanic activity is represented by basaltic lava flows, dated 65.5
140
Ma (close to the Cretaceous-Paleogene limit), 62 Ma (Danian, Paleocene), 53 Ma
7 141
(Ypresian, Eocene), and alkaline basalts, diabase and volcanic tuffs at approximately 43
142
Ma (Lutetian, Eocene).
143
3. Analytical methods
144
More than 300 core plugs and sidewall core samples from four wells were taken
145
from pervasively dolomitized and silicified intervals of the Campos Basin Pre-Salt
146
rocks (Fig. 1). Oil and salt were removed from all samples, using toluene and methanol
147
in a Soxhlet apparatus. Standard-thickness thin sections were prepared from blue epoxy
148
resin-impregnated samples for transmitted light microscopy. All thin sections were
149
examined in petrographic microscopes under uncrossed (//P) and crossed polarizers
150
(XP).
151
The paragenetic evolution of diagenetic mineral phases was examined in
152
approximately forty polished thin sections using a combination of cathodoluminescence,
153
scanning electron microscopy and microprobe analyses. Cathodoluminescence (CL)
154
was performed with a CITL Mk5-2 equipment on a Zeiss Axiocam MRC microscope.
155
Scanning electron microscopy images with secondary electrons (SEM) and
156
backscattered electrons (BSE) were obtained with a JSM-IT300 JEOL electron
157
microscope. Mineral chemical composition was verified with the support of
158
wavelength-dispersive X-ray spectroscopy (WDS) in a JEOL JXA-8230 electron
159
microprobe, and energy-dispersive spectrometry (EDS) in an OXFORD X-MaxN
160
equipment. Mineralogical mapping was also performed by QEMSCAN 650 (FEI)
161
equipment. These analyses were performed at the Petrobras Research Center (CENPES)
162
and at the Regional Center for Technological Development and Innovation (CRTI) of
163
Goiás Federal University (UFG).
8 164
Fluid inclusion (FI) petrography and microthermometry was performed on eight
165
double-polished thin sections (120 µm-thick) prepared from sidewall core samples
166
following standard procedures (Goldstein and Reynolds, 1994). These sections were
167
selected to represent the lithologies and mineral phases of the rift and sag sections (red
168
squares in Fig. 3). FIs were analyzed in macrocrystalline quartz (MQ), Sr-barite (SB),
169
saddle dolomite (SD), blocky calcite in recrystallized bioclasts (MBCB), and
170
macrocrystalline (MC), fascicular (FC) and spherulitic (CS) calcite as described in the
171
results section. A total of 203 determinations of homogenization temperature (Th) and
172
ice melting temperature (Tm ICE) of fluid inclusions hosted in syngenetic, diagenetic and
173
hydrothermal mineral phases were obtained using a Fluid Inc. USGS gas flow and a
174
Linkam THMSG600 heating and freezing stages (Table 1), under plane polarized light
175
and ultraviolet (UV; 380 nm) illumination.
176
FIs where described using the approach of Roedder (1984), where FIs where
177
classified according to its genetic type (primary, secondary, and pseudo-secondary) and
178
tied-in to the paragenetic sequence. FI types were characterized according to their
179
location, relationship to the host mineral, and consistency of visual parameters (e.g.,
180
apparent liquid/vapor ratio). Stage calibration was carried out using H2O-CO2 and pure
181
H2O-critical density synthetic inclusions. Homogenization and melting temperatures
182
were measured with a precision of ±0.1 °C. Minimum salinities were estimated using
183
the ice melting temperatures computed from NaCl-H2O system (Bodnar and Vityk,
184
1994). Salinities of inclusions displaying CaCl2-NaCl-H2O eutectic temperatures (-
185
52 °C) were estimated by combining of hydrohalite and ice melting temperatures
186
(Steele-MacInnis et al., 2011). For petroleum inclusions, API gravity was estimated
187
with a microthermometric technique, based on the oil fluorescence results (Goldstein
188
and Reynolds, 1994).
9 189
A total of 69 carbon (δ13C) and oxygen (δ18O) isotopes analyses were performed
190
on 51 samples derived from core plugs and sidewall cores of three wells (A, B and C)
191
(Table
192
microcrystalline/blocky calcite from bioclasts (MBCB) and laminites (MBCL),
193
microcrystalline/blocky dolomite from dolostones (MBDD), saddle dolomite (SD), and
194
macrocrystalline (MC), fascicular (FC) and spherulitic (CS) calcite. Forty-eight bulk
195
rock analyses were obtained from samples with more than 90% of a single carbonate
196
phase, using an IRMS Thermo Scientific Delta XL Advantage coupled to a Gasbench II
197
device in the University of São Paulo (USP) Lab. In addition, 21 samples (~100 mg) of
198
selected carbonate phases were obtained for carbon and oxygen isotopes from eight
199
polished thin sections (120 µm-thick), using a computer-monitored New Wave
200
Research MicroMill TM equipment, and a Thermo Scientific Delta V Isotope Ratio
201
Mass Spectrometer (IRMS), coupled to a Gasbench II unit in CENPES. All data were
202
reported relative to the Vienna Pee Dee Belemnite (V-PDB) international standard in
203
per mil units (‰). External precision and accuracy were checked in relation to the
204
reference material NBS 19 (TS-Limestone; Coplen et al., 2006), yielding results better
205
than ±0.09‰ for δ13C and ±0.10‰ for δ18O for both laboratories.
2;
blue
circles
The analyses of
206
87
in
Fig.
3).
The
analyzed
mineral
phases
were
Sr/86Sr ratios of thirty samples were obtained at the
207
Geochronology Lab of the University of Brasília (UnB). Roughly 1 mg of carbonate
208
powder was placed into SavillexTM (PFA tubes) and digested using an 0.5 mol/L acetic
209
acid solution in order to avoid leaching of non-carbonate phases. After centrifugation
210
the supernatant was dried in a hot plate and dissolved in nitric acid (2.9 mol/L). Sr was
211
separated from Rb in Teflon chromatographic columns filled with SrSpec (100-150 µm)
212
resin.
213
mass spectrometer in static multicollector (TIMS). The Sr ratios were normalized
87
Sr/86Sr ratios were determined using a Thermo Triton Plus thermal ionization
10 214
assuming a 88Sr/86Sr ratio of 0.1194. External reproducibility of 87Sr/86Sr ratios from the
215
NBS-987 standard yielded 0.710271 ± 0.000017 (n=26). Sr isotope analyses were
216
normalized to NBS-987 value of 0.710250.
217
4. Results
218
4.1. Petrography
219
Bioclastic grainstones and rudstones, stevensitic ooidal arenites (stevensite is a
220
smectite group clay mineral), and dolostones are the main lithologic types in the rift
221
interval (Herlinger Jr. et al., 2017; Lima and De Ros, 2019; Fig. 4A and B). Bivalves,
222
ostracods, and gastropods are the main bioclasts in the grainstones and rudstones.
223
Common diagenetic processes in these rocks include bioclast micritization,
224
neomorphism, recrystallization, replacement and dissolution. Most of the bivalves occur
225
recrystallized, and/or dissolved and filled by diagenetic cements. Cementation of inter-
226
and intraparticle, intercrystalline, moldic, vugular, channel and fracture porosity types
227
occurs mainly by drusy/rim pore-lining calcite, blocky and saddle dolomite, fibrous
228
chalcedony, microcrystalline, prismatic to coarse mosaic quartz, radiated and prismatic
229
celestine and Sr-barite, pyrite, svanbergite and bitumen. The bioclasts and the
230
stevensitic ooids are heterogeneously replaced by dolomite, calcite, microcrystalline
231
quartz, fibrous chalcedony, sulfides and svanbergite.
232
Lithologies described in the sag section include fascicular calcite crusts,
233
stevensitic claystones with calcite spherulites, intraclastic rudstones and grainstones,
234
laminites, dolostones, and cherts (Herlinger Jr. et al., 2017; Lima and De Ros, 2019; Fig.
235
4C to 4F). As for the rift bioclasts, most spherulitic and fascicular calcites are
236
recrystallized and/or replaced by silica or dolomite. The cherts consist of micro- and
237
macrocrystalline quartz, and spheroidal, fibrous chalcedony replacing syngenetic and
11 238
diagenetic constituents and filling pores. Rift lithologic types present a diagenetic
239
sequence similar to that of sag rocks. Therefore, we represent all diagenetic sequence by
240
using Figure 5, which shows the paragenetic sequence interpreted for the sag fascicular
241
calcite crusts, including 28 syngenetic, eodiagenetic, mesodiagenetic and hydrothermal
242
processes and products, schematically represented in Figure 6 (for more details in Lima
243
and De Ros, 2019).
244
Deposits of stevensite and other Mg silicates (event 1 in Fig. 5) occur throughout
245
the whole Aptian Pre-Salt succession, forming the substrate to the precipitation of
246
fascicular calcite crusts and spherulites (events 2 and 3 in Fig. 5) in a hyper-alkaline sag
247
lacustrine environment (Fig. 6A). Fascicular calcite crusts occur intercalated with
248
discontinuous levels of clay laminations, clay peloids and ooids of and siliciclastic
249
grains, such as muscovite and quartz, which were partially to totally replaced by
250
dolomite, calcite, microquartz, chalcedony and pyrite. Calcite spherulites were formed
251
as concretions, replacing and/or displacing these fine-grained deposits. Transition forms
252
between spherulitic and fascicular aggregates are common, including asymmetrical
253
spherulites, vertically elongated spherulites with lobate borders, hemi-spherulites, and
254
single or multiple fascicular aggregates nucleated on spherulites.
255
Sag lithologic types often occur partially to fully recrystallized, dolomitized,
256
silicified, cemented and/or dissolved. The Mg clay deposits occur frequently dissolved,
257
and/or replaced by silica, calcite, dolomite, pyrite or magnesite (events 4–7 in Fig. 5;
258
Fig. 6B). Lamellar magnesite replaced pseudomorphically the clay laminations in
259
subordinate amounts (event 6 in Fig. 5). Blocky dolomite, microcrystalline silica, quartz
260
and chalcedony commonly replaced the Mg clay matrix, calcite spherulites and
261
fascicular aggregates, and fill primary and secondary pores (events 8–10 in Fig. 5; Fig.
262
6B, C). Recrystallization, fracturing and pressure dissolution of fascicular and
12 263
spherulitic calcite was heterogeneous (events 11–15 in Fig. 5; Fig. 6C, D). Stylolites
264
occur
265
concentrations.
frequently associated
with
dolomitized
intervals
and
residual
pyrite
266
In some dolomitized or silicified samples, recognizing the original carbonate
267
fabrics modified by these processes is difficult, if not impossible. Dolomitization and
268
silicification are pervasive as disseminated, replacive, and filling cavities and fractures
269
associated with hydrothermal features and minerals (Figs. 7, 8). The common sag
270
replacive and pore-filling phases are cryptocrystalline to microcrystalline silica (events
271
10, 17 and 21 in Fig. 5; Fig. 7C, E); fibrous, spherulitic chalcedony (events 10 and 17 in
272
Fig. 5; Fig. 7D, E); mosaic, bladed and prismatic rims to drusiform calcite (events 12,
273
16 and 23 in Fig. 5; Figs. 7B, 8A, B, C, E, F); blocky, rims, mosaic, and saddle
274
dolomite (events 17 and 20 in Fig. 5; Figs. 7A, B, C, 8B, F); and macrocrystalline,
275
flamboyant and prismatic rims to drusiform quartz (events 21 and 22 in Fig. 5; Fig. 7D,
276
E, F).
277
Partial to total diagenetic recrystallization, dolomitization and/or silicification of
278
fascicular and spherulitic calcite are common (events 16–17 in Fig. 5; Fig. 6D, 8F). Late,
279
coarse, subhedral saddle or baroque dolomite, mega-quartz, and macrocrystalline calcite
280
(events 20, 22 and 23 in Fig. 5; Figs. 7, 8) is widely observed filling vugs and veins.
281
Saddle dolomite exhibits a characteristic xenotopic-C texture with curved crystal faces
282
and sweeping extinction (Gregg and Sibley, 1984; Davies and Smith, 2006; Fig. 7A).
283
The pervasive hydrothermal silicification (events 17, 21 and 22 in Fig 5; Fig. 6D, E, F)
284
is texturally distinct from the eodiagenetic silica crusts (event 10; Fig. 6C), which are
285
concordant
286
dolomitization. A complex evolution in pore fluid chemistry can be interpreted for the
287
hydrothermal alterations of Pre-Salt carbonate rocks.
with
depositional
bedding,
and
clearly occurred
subsequent
to
13 288
Svanbergite (SrAl3(PO4)(SO4)(OH)6) 1 to 100 µm crystals (event 18 in Fig. 5)
289
occur disseminated in microcrystalline silica and/or in chalcedony spherulites,
290
dominantly where silicification affected clay-rich protoliths. Radiated, prismatic and
291
oriented Sr-barite and celestine (event 24 in Fig. 5; Figs. 7B, F, 8A, C, D); fibro-
292
radiated zeolite (event 25 in Fig. 5; Fig. 8E); micro- to macrocrystalline, blocky and
293
prismatic sulfides (pyrite, sphalerite, chalcopyrite and galena; event 26 in Fig. 5; Fig.
294
7F); microcrystalline rutile associated with traces of native copper and zinc (event 27 in
295
Fig. 5); and massive bitumen with shrinkage cracks (event 28 in Fig. 5; Fig. 8F) are less
296
common. Sulfates, sulfides, zeolite and rutile occur filling vugular and fracture porosity,
297
often replacing dolomite, calcite and quartz. Sulfides are dominantly pyrite with lesser
298
sphalerite and minor chalcopyrite and galena. Fluorite and dickite were also described
299
in other wells of the area (Herlinger Jr. et al., 2017), and dawsonite was identified in the
300
XRD analyses (Lima and De Ros, 2019).
301
The pore types that occur in the sag lithologic types are growth-framework,
302
inter-crystalline aggregates, from Mg-clay dissolution, intercrystalline and intra-
303
crystalline aggregates, vugular, and fracture (Lima and De Ros, 2019). Important
304
fracturing and dissolution events occurred, which largely modified the pre-existing pore
305
systems (events 7, 13, 14, 19 and 21 in Fig. 5; Fig. 6). The hydrothermal alterations
306
were related to intense fracturing and dissolution processes (events 14, 19 and 21; Fig.
307
6D, E and F). The late hydrothermal saddle dolomite, macrocrystalline quartz and
308
calcite, sulfates, sulfides, zeolite, svanbergite, rutile, and bitumen filled growth
309
framework, and mostly fracture and dissolution pores (events 18–28 in Fig 5; Fig. 6D, E,
310
F). A later overprint of calcite-quartz-sulfides veins and breccias is also observed.
311
In general, microcrystalline (MD), blocky (BD) and saddle dolomite (SD)
312
present higher concentrations of Na2O, Al2O3, SiO2, MnO and FeO compared to
14 313
fascicular (FC), spherulitic (CS) and macrocrystalline calcite (MC) (WDS analysis; see
314
Supplementary material). SD shows higher values of SiO2 and lower MnO and FeO
315
than other dolomite types. Similarly, hydrothermal MC also has lower concentrations of
316
SiO2, SrO and BaO than FC and CS. In addition, hydrothermal sphalerite (Sp) shows
317
high Fe and Pb and low As, Co, Cu and Ni contents in WDS analyses (Supplementary
318
material). Hydrothermal euhedral pyrite (Py) filling secondary porosity also occurs with
319
high Pb and low As, Co, Cu, Zn and Ni concentrations (WDS). Hydrothermal prismatic
320
Sr-barite (SB) presents SrO content ranging from 2% to 16%, with an average of 6.7%,
321
high TiO2 and low CaO and FeO concentrations (WDS).
322
4.2. Fluid inclusion analyses
323
Aqueous and petroleum fluid inclusions (FIs) hosted in syngenetic and
324
diagenetic/hydrothermal constituents (Fig. 9) were analyzed in the main rift and sag
325
lithologic types (Figs. 4, 7, 8), such as fascicular calcite crusts, stevensitic claystones
326
with calcite spherulites, intraclastic grainstones, bioclastic rudstones and cherts. The
327
phases analyzed were recrystallized bioclasts, fascicular and spherulitic calcite, and
328
diagenetic/hydrothermal macrocrystalline calcite, saddle dolomite, mega-quartz, and
329
prismatic Sr-barite, replacing pre-existing constituents, filling fractures and dissolution
330
porosity. Fluid inclusions (both aqueous and petroleum inclusions) are trapped in
331
recrystallized fascicular and spherulitic calcite displaying randomly distribution in areas
332
with clear mineral substitution. On the other hand, the preserved carbonate elements do
333
not host petroleum inclusions. Therefore, these inclusions in the recrystallized
334
carbonates
335
microthermometric results are displayed in Table 1.
were
interpreted
as
primary.
Petrographic
observations
and
15 336
All aqueous inclusions are two-phase (liquid plus vapor) at room temperature
337
without large variation in the vapor fraction, with exception of some primary inclusions
338
in saddle dolomite and macrocrystalline calcite displaying necking-down features,
339
which were disregarded from subsequent interpretations. Six high Th results from
340
saddle dolomite (140, 144, 148 and 150 °C) and Sr-barite (144 and 152 °C) may
341
represent stretched inclusions or unrecognized necking after a phase change (Goldstein,
342
2001).
343
Analyzed aqueous FIs were hosted in the following mineral phases:
344
recrystallized fascicular calcite (n=2), mega-quartz (n=34), prismatic Sr-barite (n=20),
345
macrocrystalline calcite (n=36), and saddle dolomite (n=34). The majority of the FIs
346
observed were primary and pseudo-secondary aqueous and petroleum types (78%), with
347
subordinate secondary inclusions (22%; Table 1). Aqueous secondary FIs are more
348
significant in macrocrystalline quartz (38%). The salinities of aqueous inclusions varied
349
between 12 and 26 wt. % eq. NaCl (Table 1).
350
Petroleum inclusion from sag samples correspond predominantly to white-blue-
351
fluorescent inclusions (n=82) and less yellow-fluorescent inclusions (n=16). Most are
352
hosted in the recrystallized fascicular and spherulitic calcite, and in the macrocrystalline
353
quartz, and less in macrocrystalline calcite, Sr-barite, and saddle dolomite (Table 1).
354
Petroleum FIs in rift samples consist of mostly white-blue-fluorescent (n=20) in
355
recrystallized calcite bioclasts, and less of yellow-fluorescent (n=3) in saddle dolomite,
356
macrocrystalline quartz and Sr-barite. Estimated API gravities range from 32 to 35° for
357
the yellow-fluorescent petroleum inclusions, and between 40 and 50° for the white-
358
blue-fluorescent inclusions. The characteristics of the FIs in specific analyzed
359
constituents can be summarized as follows:
16 360
Recrystallized calcite bioclasts (CB): 20 analyses of FIs in blocky calcite in
361
recrystallized calcite bioclasts (CB) from a rift rudstone in well B (Table 1; Figs. 3,
362
10B). Only primary petroleum FIs with white fluorescence (~45 to 50° API) were
363
observed. The Th values vary from 80 to 108.6 °C, with mean and median of 94.8 and
364
96.9 °C.
365
Recrystallized fascicular calcite (FC): 22 FI analyses in a sag recrystallized
366
syngenetic fascicular calcite (FC) sample from well B (Table 1; Figs. 3, 10C). Two
367
aqueous FIs present Th of 92 and 93 °C, and salinity of 23.9 and 24.3 wt. % eq. NaCl
368
(Fig. 11). White to blue fluorescent petroleum FIs (~40 to 50° API) have Th ranging
369
from 83 to 111 °C.
370
Recrystallized spherulitic calcite (CS): 38 FIs were analyzed in recrystallized
371
eodiagenetic calcite spherulite (CS) in a sag stevensitic claystone from well A (Table 1;
372
Figs. 3, 10D, E). Both primary and secondary petroleum inclusions with white
373
fluorescence color (~45 to 50° API) were hosted in the CS. Th of 18 primary petroleum
374
inclusions varies from 63 to 112.2 °C, with mean and mode of 83.5 °C (Fig. 10D). Th of
375
20 secondary petroleum inclusions ranges from 63 to 97.4 °C, with mean of 75.7 and
376
median of 72.1 °C.
377
Macrocrystalline calcite (MC): 36 primary and secondary aqueous and
378
petroleum FIs were measured in macrocrystalline calcite (MC) filling dissolution
379
porosity of three sag samples from well C (Table 1; Figs. 3, 12G, H). The 22 primary
380
and pseudo-secondary aqueous FIs present Th varying from 101 to 130 °C, with mean
381
of 114.3 and mode of 111.5 °C. Salinity of primary FIs ranges from 15.1 to 22.9 wt. %
382
eq. NaCl (Fig. 11A). Th of 5 secondary aqueous FIs varies from 99 to 126 °C, with
383
mean of 107 and mode of 101 °C (Fig. 12H). Salinity of secondary aqueous FIs ranges
384
from 18.8 to 22.9 wt. % eq. NaCl (Fig. 11B). Yellow-fluorescent (~30 to 35 °API)
17 385
primary inclusions have Th value of 100 °C. Th in 5 secondary petroleum FIs range
386
from 69 to 124 °C, with mean of 93.8 and median of 81 °C.
387
Macrocrystalline quartz (MQ): 58 FIs were observed in pore-filling mega-quartz
388
(MQ) in a silicified intraclastic grainstone, a FC crust, a chert and a stevensitic
389
claystone with CS from the sag phase in wells A and C (Table 1; Figs. 3, 12E, F). 16
390
primary/pseudo-secondary aqueous FIs display Th from 105 to 139 °C, with mean of
391
122.7 and mode of 132.5 °C (Fig. 12E). Salinity of these inclusions ranges from 12.9 to
392
23.7 wt. % eq. NaCl (Fig. 11A). Th of secondary aqueous FIs varies from 105 to 134 °C,
393
with mean of 116.6 and mode of 118.5 °C (Fig. 12F). Salinity of secondary aqueous
394
inclusions ranges from 19.9 to 23.9 wt. % eq. NaCl (Fig. 11B). Th of 26 white to yellow
395
fluorescent petroleum inclusions (~32 to 50° API) varies from 65 to 100.9 °C, with
396
mean of 76.3 and mode of 73.1 °C (Fig. 12E).
397
Prismatic Sr-barite (SB): 20 FIs analyses in prismatic Sr-barite (SB) cement
398
filling secondary porosity in a sag silicified intraclastic grainstone in well C (Table 1;
399
Figs. 3, 12C, D). Th of 11 primary/pseudo-secondary aqueous FIs ranges from 122 to
400
152 °C, with mean of 131.1 and mode of 127.5 °C (Fig. 12C). Th of 3 secondary
401
aqueous FIs varies from 122 to 124 °C (Fig. 12D). Salinity ranges from 16.7 to 22.1
402
wt. % eq. NaCl in 10 primary aqueous (Fig. 11A), and 20.2 to 21.8 wt. % eq. NaCl in 2
403
secondary aqueous FIs (Fig. 11B). Both primary and secondary petroleum FIs show
404
yellow fluorescence colors (~32 to 34° API). Only one primary yellow-fluorescent
405
petroleum inclusion yielded a Th value of 83 °C (Fig. 12C). Secondary petroleum FIs
406
have Th ranging from 100 to 119 °C with mean of 107.8 and mode of 106 °C (Fig. 12B).
407
Saddle dolomite (SD): 34 FIs in pore-filling saddle dolomite (SD) in sag FC
408
crusts and rift bioclastic rudstones from well C (Table 1; Figs. 3, 12I). Th measured in
409
20 primary aqueous FIs in zoned saddle dolomite varies from 105 to 150 °C with mean
18 410
of 132.8 and mode of 134.5 °C (Fig. 12I). Salinity of 29 primary aqueous FIs in SD
411
ranges from 16.9 to 26.1 wt. % eq. NaCl (Fig. 11A). Th obtained in 3 yellow-
412
fluorescent primary petroleum inclusions varies from 79 to 89 °C (Fig. 12I).
413
The salinities of inclusions hosted in recrystallized FC are close to the upper
414
limit for pore-filling hydrothermal minerals (Fig. 11A), with 97% higher than 15 wt. %
415
eq. NaCl. In the rift interval, salinities range from 17.3 to 22.7 wt. % eq. NaCl. The final
416
melting temperature of aqueous FIs hosted in SD, MC, MQ and SB ranges of -9 to -
417
25.3 °C, indicating NaCl-dominant compositions comparable to CaCl2 (e.g., Roedder
418
and Bodnar, 1997; Wilkinson, 2010).
419
Th of white-blue-fluorescent primary petroleum inclusions hosted in the
420
recrystallized CB, CS and FC ranges from 63 to 112.2 °C (Table 1; Fig. 10). Th of
421
yellow-fluorescent inclusions hosted in MC, MQ, SB and SD varies between 69 and
422
124 °C. Some yellow-fluorescent oil inclusions are interpreted to have been emplaced at
423
105 to 108 °C in MC and at 122 to 130 °C in SB in the presence of hypersaline aqueous
424
fluids. Mixed, co-genetic oil-aqueous FIs with Th ranging from 79 to 124 °C were
425
observed in fascicular calcite crusts, silicified intraclastic grainstone and bioclastic
426
rudstone (Table 1), indicating a common origin from a single fluid. Aqueous inclusions
427
coexist with yellow-fluorescent petroleum inclusions in MC, SB and SD.
428
4.3. Carbon and oxygen isotopes
429
The range of δ13C values for all analyzed carbonates is quite narrow, varying
430
between -1.88 and 2.43‰ (Table 2; Fig. 13A, B). The range of δ13C values for the
431
seven different carbonate phases are: 1) 0.45 to 1.43‰ for microcrystalline/blocky
432
calcite
433
microcrystalline/blocky calcite from laminites (MBCL); 3) 1.16 to 1.44‰ for
from
bioclasts
(MBCB);
2)
0.93
to
1.52‰
for
eodiagenetic
19 434
mesodiagenetic microcrystalline/blocky dolomite from dolostones (MBDD); 4) 0.31 to
435
2.43‰ for recrystallized eodiagenetic calcite spherulites (CS); 5) 0.05 to 1.98‰ for
436
recrystallized syngenetic fascicular calcite (FC); 6) -0.27 to 1.28‰ for diagenetic
437
macrocrystalline calcite (MC); and, 7) -1.88 to 1.03‰ for saddle dolomite (SD).
438
In contrast to the δ13C values, the δ18O values from the analyzed carbonates
439
show a wider variation range (from -9.74 to 1.74‰) (Table 2; Fig. 13A, B). The range
440
of δ18O values from the seven different carbonate phases are: 1) -1.18 to 1.04‰ for
441
MBCB; 2) -0.99 to 0.63‰ for MBCL; 3) -1.51 to 0.35‰ for MBDD; 4) -2.11 to 1.74‰
442
for CS; 5) -5.70 to -1.93‰ for FC; 6) -9.74 to -4.92‰ for MC; and, 7) -9.49 to -6.35‰
443
for SD. Samples analyzed in this study may be arranged into two main groups based on
444
the δ18O values: samples (MC and SD) with more depleted oxygen isotope values (<-
445
4.92‰); and samples (MBCB, MBCL, MBDD and CS) with more enriched oxygen
446
isotope values (>-2.11‰) (Fig. 13A). The δ18O values of FC samples (-5.70 to -1.93‰)
447
fall in between these two groups (Fig. 13A). The comparison of Fig. 13A and 13B with
448
13C and 13D show that both δ13C and δ18O values from different mineral phases are not
449
biased by different wells. Likewise, the distribution of samples against stratigraphy (Fig.
450
3) shows that sampling for isotope analyses are not biased against stratigraphy.
451 452
4.4. Strontium isotopes The analyzed samples show a wide range of
87
Sr/86Sr values (0.71100 to
453
0.71394; Table 3; Fig. 13). They may be separated into two main groups: those with Sr
454
isotope ratios above 0.71299 (MBDD, CS, FC, MC) and those below 0.71226 (MC, SD,
455
SB). As shown in Table 3, the Sr isotope values have a narrow range within each group
456
of samples. For instance, SD values range between 0.71100 and 0.71124, while CS
457
samples range between 0.71299 and 0.71394. As in the δ13C and δ18O results, 87Sr/86Sr
20 458
ratios from different mineral phases are not biased against stratigraphy (Fig. 3) or by
459
different wells (Fig. 13).
460
5. Discussion
461
5.1. Fluid temperature and salinity
462
There is no correlation between the salinity and temperature of fluid inclusions
463
(Fig. 14). Except for few FIs in macrocrystalline quartz and calcite, all other phases
464
exhibit similar salinity range. In contrast, the average Th of saddle dolomite and Sr-
465
barite inclusions are higher than those measured in the mega-quartz and calcite. Based
466
on these observations, the following points may be considered: i) the salinity variations
467
do not indicate mixing of fluids, but are related to cooling and/or neutralization by fluid-
468
rock interaction; ii) the formation of mega-quartz, saddle dolomite, macrocrystalline
469
calcite and Sr-barite, which are associated with higher temperature fluids, postdated the
470
formation of the other phases. This further indicates that heat was not pervasive in the
471
system, but rather concentrated along structures such as fractures that acted as conduits
472
for the hot fluids.
473
The variability of the Th values in aqueous inclusions may be related to post-
474
entrapment modification of FIs at variable high temperatures (King and Goldstein,
475
2018). For instance, the necking-down process occurs most rapidly at higher
476
temperatures, often producing Th values without any relation to the original
477
temperatures (Roedder, 1984; Bodnar et al., 1985; Goldstein, 2001; Bodnar, 2003). The
478
similar Th values in macrocrystalline calcite and quartz indicate that there were no
479
significant post-entrapment changes in FIs. Another relevant indication of preservation
480
of the integrity of the FIs is the robust petrographic control evidenced by the relatively
481
narrow range of Th values of individual fluid inclusions, generally much narrower than
21 482
within the entire sample, or between different samples (e.g., Wilkinson, 2010).
483
Therefore, the wider dispersion in Th values observed in hydrothermal saddle dolomite
484
is probably representing different generations of FIs trapped in the zoned crystals (Fig.
485
9D).
486
The Th values of aqueous FIs hosted in minerals precipitated or recrystallized by
487
hydrothermal fluids show a relatively narrow range, varying from to 92 and 152 °C
488
(Table 1; Fig. 14), in which 95% of the results range from 100 to 140 °C. In general,
489
there is a remarkable Th overlapping among all analyzed phases, with slightly higher
490
mean values for Sr-barite and saddle dolomite. The FI data indicate that hydrothermal
491
saddle dolomite, Sr-barite, macrocrystalline calcite, and mega-quartz precipitated over a
492
similar temperature range. The only two analyses performed in aqueous FIs of
493
recrystallized fascicular calcite yielded Th near the lower limits of the hydrothermal
494
phases (Table 1; Fig. 14). Poros et al. (2017) verified a Th variation between 95 to
495
115 °C in inclusions hosted in dolomite cement, 95 to 125 °C in silica (chert and
496
chalcedony) replacing carbonate, and 105 to 125 °C in pore-filling mega-quartz in the
497
Pre-Salt of Kwanza Basin, Angola. Although those mineral phases were interpreted to
498
have been formed by a hydrothermal system in a deep-burial condition, the Th values in
499
Kwanza Basin Pre-Salt reservoirs are somewhat lower than those obtained in Campos
500
Basin.
501
All analyzed phases in the study area contain abundant oil and/or co-genetic oil-
502
aqueous inclusions, which is indicative of atypical petroleum generation and migration
503
(sensu Magoon and Dow, 1994) associated with percolation of hydrothermal fluids
504
responsible for the recrystallization and precipitation of the host mineral phases. When a
505
sedimentary sequence is percolated by hot fluids, the organic matter of host rocks can
506
be thermally altered, generating petroleum through hydrothermal or “forced” maturation
22 507
(Ilchik et al., 1986; Anderson, 1991; Davies and Smith, 2006). The liquid petroleum
508
droplets produced by this mechanism are transported by the convecting fluids
509
(Schoenherr et al., 2007), and can be either trapped as petroleum FIs in hydrothermal
510
minerals, or devolatilized in the pore space as bitumen. This type of natural hydrous
511
pyrolysis is not efficient enough to fill a reservoir or an oil field (Davies and Smith,
512
2006) but may, however, generate oil or co-genetic oil-aqueous FIs and solid bitumen
513
(Simoneit, 1990; Simoneit, 2018, and references therein) comparable to those described
514
here, which may be associated to hydrothermal dolomitization and ore deposits
515
described in diverse geotectonic settings (Disnar, 1996; Wilson and Zentilli, 1999;
516
Parnell et al., 2003; Bertrand et al., 2003; Neilson and Oxtoby, 2008; Suchý et al., 2010;
517
Greenwood et al., 2013 and references therein; Ostendorf et al., 2015).
518
With progressive burial, ‘normal’ background mesodiagenetic processes
519
proceeded in Campos Basin Pre-Salt independently of hydrothermal events (Fig. 5).
520
Thus, it should be noted that part of the petroleum inclusions hosted in the recrystallized
521
calcite bioclasts and spherulites (Th = 80–111 °C), in which no aqueous inclusions were
522
observed, may have been trapped during the migration of oil generated by the
523
conventional petroleum system. According to Mello et al. (1994), conventional burial
524
petroleum generation, migration and accumulation started in Campos Basin only in the
525
Miocene, being thus subsequent to the atypical hydrothermal petroleum migration. This
526
conventional petroleum system was responsible for the large volume of oil and gas that
527
fills the Pre-Salt reservoirs.
528
Salinities are consistently high in aqueous inclusions in the studied hydrothermal
529
phases (12 to 26 wt. % eq. NaCl; Table 1), suggesting that they were formed from
530
hypersaline brines. In this context, significant Th variation within a narrow salinity
531
range indicates temperature variation within the hydrothermal system during mineral
23 532
precipitation (Fig. 14). Nevertheless, no correlation was observed between Th and
533
salinity data. This, together with homogeneous high salinity values, indicates that there
534
was no fluid mixing in the studied hydrothermal system, and that the minerals
535
precipitated from a single hydrothermal system. Rare salinities below 15 wt. % eq.
536
NaCl from aqueous FIs hosted in macrocrystalline calcite and quartz (Table 1; Fig. 14)
537
could indicate some incipient isothermal mixing between the predominant hydrothermal
538
fluid (approximately 20–21 wt. % eq. NaCl) with a moderate-salinity fluid
539
(approximately 13 wt. % eq. NaCl). However, since 97% of the salinity results are
540
above 15 wt. % eq. NaCl, this relatively lower values may be related to salinity
541
variation in recharge area (Khaska et al., 2013).
542
The Th data in the range 92 to 152 °C with salinities of 13 to 26 wt. % eq. NaCl
543
of Campos Basin Pre-Salt are very similar to the majority of hydrothermal systems
544
observed in the hydrothermal dolomites (HTD), carbonate-hosted Pb-Zn Mississippi
545
Valley (MVT) and Irish-type deposits (Davies and Smith, 2006; Leach et al., 2001;
546
Wilkinson, 2001; Paradis et al., 2007; Wilkinson, 2010).
547
5.2. Carbon, oxygen and strontium isotopes
548
The δ13C values of host rock mineral phases (MBCB, MBCL, MBDD, FC and
549
CS) and hydrothermal mineral phases (SD and MC) present a very narrow variation
550
(between -1.88 and 2.43‰; Fig. 13, Table 2). These values are similar to those
551
previously reported from Pre-Salt carbonates from Campos, Santos and Kwanza Basins.
552
δ13C bulk-rock data from the cuttings of well CP-5 in Campos Basin vary between -0.38
553
and 2.71‰ for the Macabu Formation, and between -0.88 and 1.35‰ for the Coqueiros
554
Formation (Dias, 1998; Rodrigues, 2005; Muniz and Bosence, 2015). The δ13C bulk-
555
rock data (-1.52 to 3.20‰) of spherulite, shrub, laminate, calcarenite and calcirudite
24 556
lithofacies from core samples of the sag Barra Velha Formation from Santos Basin
557
(Farias et al., 2019) are similar to reported values of lithofacies from the Macabu
558
Formation (Herlinger Jr. et al., 2017; Lima and De Ros, 2019).
559
Syn-rift and sag intervals from Kwanza Basin have lithofacies and mineral
560
phases similar to those described for the rift (Coqueiros Formation) and sag (Macabu
561
Formation) intervals of the Campos Basin (Fig. 2). The δ13C values of spherulitic and
562
shrubby calcite and early dolomite from the sag interval of Kwanza Basin vary between
563
-4.3 and 3‰, whereas neomorphic calcite and calcite cement filling mollusk molds and
564
interparticle pores from syn-rift interval of Kwanza Basin usually display δ13C between
565
-3.0 and 2.0‰ (Saller et al., 2016; Sabato Ceraldi and Green, 2016; Fig. 13). Although
566
Sabato Ceraldi and Green (2016) present δ13C and δ18O values from rift and sag Pre-Salt
567
sections from Kwanza Basin, it is not possible to identify the carbonate analyzed. The
568
lithofacies, mineralogy and δ13C data of Campos, Santos and Kwanza basins suggest
569
that the depositional and diagenetic processes were quite similar within the vast Aptian
570
lacustrine system, indicating either a single water body or hydrogeological connected
571
lakes.
572
Data from the present study show that δ18O values of host rock mineral phases
573
(MBCB, MBCL, MBDD and CS) vary between -2.11 and 1.74‰, except for FC that
574
has values raging between -5.70 and -1.93‰ (Fig. 13, Table 2). Similar pattern was
575
obtained by Sabato Ceraldi and Green (2016). In contrast, hydrothermal mineral phases
576
(SD and MC) present much lower isotope values (-9.74 to -4.92‰; Fig. 13, Table 2),
577
suggesting either differences in the temperature of the fluid and/or in their oxygen
578
isotopic composition (δ18Ow). In addition, δ18O values of FC samples suggest
579
recrystallization with temperature and/or δ18Ow different from both other host rock
580
mineral phases and hydrothermal mineral phases. The comparison between the δ18O of
25 581
carbonates from the Pre-Salt and others formed elsewhere is not straightforward,
582
because their isotopic composition depends on the temperature, carbonate phase and
583
fluid δ18Ow (Sharp, 2007). Equations (1) and (2) show the relationship between δ18Ocarb,
584
δ18Ow and temperature for calcite (O’Neil et al., 1969) and dolomite (Matthews and
585
Katz, 1977), respectively:
586
10 ∝ = 2.78 ×
587
10 ∝ = 3.06 ×
− 2.89 (1) − 3.24
(2)
588
Where ( 10 ∝ ) is approximately equal to δ18Ocarb - δ18Ow, and T is the
589
temperature in Kelvin. These equations were used to calculate the δ18Ow of fluid in
590
equilibrium with each mineral phase (FC, CS, SD and MC) based on Th obtained from
591
aqueous and petroleum FIs, similar to the approach by King and Goldstein (2018). The
592
δ18Ow VPDB (Vienna Pee Dee Belemnite) was then converted to δ18Ow VSMOW
593
(Vienna Standard Mean Ocean Water) by using the equation 3 (O’Neil et al., 1969):
594
δ# O% &'()* = 1.03086 × δ# O% &+,- + 30.86 (3)
595
Calculated δ18Ow for FC samples vary between 5.4 and 12.4‰ (Table 4) based
596
on the maximum and minimum measured δ18Ocarb (-5.7 and -1.93‰) and Th derived
597
from petroleum FIs (83–111 °C). A variation of 6.5 to 10.4‰ was calculated for FC
598
δ18Ow using just two measurements of Th from aqueous FIs, both of 92.2 °C. Calculated
599
δ18Ow of CS samples vary between 6.3 and 16.3‰ based on the maximum and
600
minimum measured δ18Ocarb (-2.11 and 1.74‰) and Th derived from petroleum FIs (63–
601
112.2 °C). Those values calculated from Tho suggest that δ18Ow was the oxygen isotope
602
composition of pore water during the recrystallization of both mineral phases (FC and
603
CS). These high δ18Ow values are similar to those calculated for the isotopic
604
composition of fluids in equilibrium with shrub and spherulite lithofacies from Well
26 605
SB-2 of the Barra Velha Formation, Santos Basin (Farias et al., 2019). Based on oxygen
606
isotope compositions and clumped temperatures, the recalculated δ18Ow in equilibrium
607
with carbonates from these lithofacies vary between 5.0 to 10.7‰. These highly
608
enriched δ18Ow are comparable to the isotopic composition of residual evaporative
609
brines (Epstein and Mayeda, 1953) and/or of fluids resulted from extensive rock-water
610
interaction with basinal materials (King and Goldstein, 2018). Therefore, we suggest
611
that those values represent δ18Ow modified from residual evaporative brines during
612
burial, in which δ18Ow is the product of extensive rock-water interaction and
613
recrystallization. Considering this scenario, the FIs in spherulites and shrubs were
614
trapped during the recrystallization of these constituents during advanced burial. Some
615
FIs in these phases may, however, have been inherited from earlier stages of
616
neomorphism/recrystallization.
617
The oxygen isotopic composition of aqueous fluids (δ18Ow) in equilibrium with
618
MC are calculated between 3.3 and 11.1‰ (Table 4), based on the maximum and
619
minimum measured δ18Ocarb (-9.74 and -4.92‰) and on Th derived from aqueous FIs
620
(101−130 °C). These values are similar to those calculated for the aqueous fluids in
621
equilibrium with SD (δ18Ow between 2.3 and 10.0‰), which are based on aqueous FIs
622
in the range of 105 to 150 °C and maximum and minimum measured δ18Ocarb between -
623
9.49 and -6.35‰. These results indicate that MC and SD were formed from the same
624
hydrothermal fluid.
625
The
87
Sr/86Sr ratios from of host rock mineral phases (FC, CS and MBDD –
626
from 0.71299 to 0.71394) are significantly higher than those from hydrothermal mineral
627
phases (SB, SD and MC – from 0.71100 to 0.71226) (Fig. 13, Table 3). Reported
628
87
629
2017), which is the same well labeled as well A in this study, were obtained by leaching
Sr/86Sr values for the Macabu Formation (0.712991−0.713365) in well C (Tedeschi,
27 630
carbonate phases with weak acid, without differentiation of carbonate phases or micro-
631
sampling. The results presented here are also similar to previous data from other
632
Brazilian Pre-Salt wells at Campos and Santos basins (Dias, 1998; Pietzsch et al., 2018;
633
Farias et al., 2019; Fig. 13F).
634
The
87
Sr/86Sr ratios of carbonate samples from the rift Coqueiros and sag
635
Macabu formations range from 0.7117 to 0.7118 (n=2) and 0.7128 to 0.7137 (n=2)
636
(Dias, 1998). The
637
spherulite, shrub, laminate, calcarenite and calcirudite, from Santos Basin SB-2 well
638
vary from 0.7130 to 0.7138 (Farias et al., 2019). In addition, the 87Sr/86Sr ratios of 194
639
sidewall samples from well S10 in Santos Basin vary between 0.7105 and 0.7140
640
(Pietzsch et al., 2018). The isotope ratios from the rift Itapema Formation (Jiquiá local
641
stage) were usually lower than 0.7115, while the ratios from the sag Barra Velha
642
Formation (Alagoas local stage) were roughly between 0.7120 and 0.7140. These values
643
agree with five diagenetically modified ostracod samples analyzed by Tedeschi (2017),
644
three of which are from Hourcqia sp. valves (Itapema Formation;
645
0.711553 to 0.712012) and two are from Pattersoncypris ssp. valves (Barra Velha
646
Formation;
647
isotopes and 87Sr/86Sr data from different studies, depositional and diagenetic processes
648
were similar over a regional scale in Campos and Santos Basins, suggesting a single
649
vast lake or lakes that were hydrogeologically connected. Additionally, these data
650
establish constraints on
651
comparison with the hydrothermal mineral phases.
652
The
87
Sr/86Sr ratios from 43 samples from the sag phase, including
87
Sr/86Sr ratios from
87
Sr/86Sr ratios from 0.712756 to 0.713427). As indicated by the stable
87
87
Sr/86Sr ratios at a regional scale that can be used for
Sr/86Sr ratio of Pre-Salt carbonates indicate a strong interaction with 87
Sr/86Sr ratios,
653
continental crust materials, being significantly higher than marine
654
specifically than seawater values from Early Cretaceous (from 0.707074 to 0.707541;
28 655
Jenkyns et al., 1995; McArthur et al., 2012; Bodin et al., 2015; Yamamoto et al., 2013;
656
Ando, 2015) as well as that of Phanerozoic seawater (<0.7095; McArthur et al., 2012;
657
Fig. 13F). Based on Sr isotopes, Pietzsch et al. (2018) presented a geological model to
658
discriminate the Sr source of lacustrine rift and sag sections in the Santos Basin.
659
According to this model, Sr isotopes of carbonate rocks from sag Barra Velha
660
Formation were controlled by groundwater with long residence time of interaction with
661
felsic basement rocks. Although one may argue that leaching of waters through felsic
662
basement was not enough to deliver ions (Mg2+, Ca2+ and others) to develop the Barra
663
Velha Formation, a mass balance using waters from the basement has not been
664
performed. Therefore, although this process seems to be feasible at local scale, as
665
pointed out by the geochemical models of Teboul et al. (2016), the isotopic and
666
elemental aspects of the huge Pre-Salt volumes still need to be explained in basinal
667
scale.
668
Otherwise, coquinas from the Itapema Formation would have been deposited in
669
a rift lacustrine environment with higher fluvial recharge, and more variable catchment
670
area and subsurface recharge. This hydrological model for the Itapema Formation would
671
allow mixing of groundwater recharge with long residence time with waters derived
672
from Cabiúnas volcanic and volcaniclastic rocks and Piçarras Formation siliciclastic
673
sediments. Considering the 87Sr/86Sr ratios obtained from the Campos Basin, we would
674
consider the same hydrological patterns to explain the differences between
675
ratios from the rift Coqueiros and the sag Macabu formations.
676
The
87
Sr/86Sr
87
Sr/86Sr values of SD, MC and SB overlap each other (Fig. 13, Table 3).
677
Although the petrographic sequence shows minor dissolution of SD followed by
678
formation of MC and SB, there is not a definitive petrographic evidence that these
679
mineral phases are derived from different fluids (Figs. 5, 8). The same 87Sr/86Sr values
29 680
and the absence of petrographic evidence suggest that SD, MC and SB were formed
681
from the same fluid. In addition, strontium isotope composition of carbonate
682
hydrothermal phases (SD and MC) shows respectively positive and negative correlation
683
with δ13C and δ18O values, which also suggest a same trend of evolution from the same
684
fluid (Fig. 13). The relationship between δ13C and
685
observed in hydrothermal baroque dolomite from Cambrian-Ordovician Arbuckle
686
Group in USA midcontinent, whereas the relationship between δ18O and 87Sr/86Sr ratios
687
is the same (King and Goldstein, 2018). These observations suggest the same strontium
688
source for all these three mineral phases. The negative correlation between δ18O and
689
87
690
different rocks or that hydrothermal fluids at higher temperature and/or with lower
691
δ18Ow would promote stronger host rock dissolution, becoming more radiogenic.
692
Likewise, the positive correlation between
693
increased host rock dissolution by the hydrothermal fluids. According to this
694
interpretation, strong dissolution of basement and rift rocks precede the formation of
695
these hydrothermal mineral phases (SD, MC and SB).
696
87
Sr/86Sr is the opposite of those
Sr/86Sr (Fig. 13) suggests either that different set of fluids were in equilibrium with
87
Sr/86Sr and δ13C would be related to
5.3. Hydrothermal alteration of carbonate rocks
697
The term hydrothermal alteration was originally applied to situations with the
698
presence of hot waters and ore deposits associated with magmatic activities (e.g.,
699
Gilbert, 1875; Morey and Niggli, 1913; Holmes, 1928; Stearns et al., 1935). More
700
recently, hydrothermal alteration was defined as a replacement of the original minerals
701
by hydrothermal fluids that deliver reactants and remove aqueous reaction products
702
(Utada, 1980; Henley and Ellis, 1983; Inoue, 1995; Reed, 1997). Hydrothermal fluids
703
represent aqueous solutions of different compositions and origins (magmatic,
30 704
metamorphic, groundwater, or meteoric waters), which are hotter than the wall-rock
705
(Skinner, 1997).
706
Hydrothermal alterations represent the evidence of a geothermal anomaly,
707
requiring both a mechanism (heat source) and conduits for the flow of fluids (e.g. deep
708
fault systems). Hydrothermal systems are characterized by focusing of fluids through
709
fault systems at very high flow rates, in a transient or episodic regime (Davies, 2004).
710
Such focalization of fluids and their interaction with the host rocks are linked to the
711
origin of carbonate-hosted base metal deposits (e.g., the Pb-Zn Mississippi Valley type
712
deposits; Leach, 1994) and may also modify significantly the original properties of
713
some hydrocarbon reservoirs (e.g., Smith, 2006; Zhu et al., 2015). The so-called
714
hydrothermal petroleum is generated in volcanic and geothermal marine and continental
715
environments, and sampled from submarine and sublacustrine oil seeps (vent fluids),
716
mud volcanoes, hot springs, chimneys, fumaroles and other surface manifestations
717
(Ventura et al., 2012; Simoneit, 2018).
718
Despite several studies on hydrothermal ore deposits, the characterization and
719
definition of hydrothermal processes in sedimentary successions are still a matter of
720
heated discussion (e.g., Machel and Lonnee, 2002; Davies, 2004; Davies and Smith,
721
2006). White (1957) defined hydrothermal processes solely based on a temperature
722
contrast of 5 °C or more between the inflowing aqueous solutions and the wall-rocks,
723
disregarding any temperature limits or source of the fluids. Similarly, Machel and
724
Lonnee (2002) suggested that minerals formed in sedimentary rocks should be
725
considered hydrothermal minerals only if precipitated from fluids with temperatures at
726
least 5 to 10 °C hotter than the host rock temperature, disregarding the fluid sources and
727
the pathways. More loosely, Davies and Smith (2006) defined hydrothermal fluids as
728
those arising to the surface at temperatures higher than those of the depositional
31 729
environment, or that are introduced into the host rocks at a higher temperature than the
730
host. Following such line of definition, sedimentary ore-forming hydrothermal systems
731
may be either depositional (syngenetic) or diagenetic (epigenetic), depending on the
732
time, depth, geometry and mechanism of the flow of fluids. In still other lines of
733
definition, the minerals are hydrothermal if considered “unusual”, or “exotic” in relation
734
to the common carbonate rock mineralogy, including saddle dolomite, fluorite, barite,
735
anhydrite, dickite, sphalerite, and pyrite (Neilson and Oxtoby, 2008).
736
As reliable determination of paleotemperature contrasts in the order of only 5 to
737
10 °C is impossible in most cases, and as minerals that may be considered “exotic” by
738
some authors are widespread in sedimentary successions, or simply do not occur in
739
several recognized hydrothermal systems, we opted for a more practical and operational
740
definition for hydrothermal processes and products. This definition is based on the
741
focusing of the flow of relatively heated fluids coming from below through faults,
742
fractures, unconformities and similar conduits, and on the consequent concentration of
743
mineral precipitation and/or dissolution in the proximity of such conduits. Therefore, in
744
this sense, temperature and/or compositional requirements are not mandatory for the
745
definition of hydrothermal alterations.
746
5.3.1. Hydrothermal dolomitization
747
Hydrothermal dolomitization was defined by Davies and Smith (2006) as
748
occurring under subsurface conditions, commonly at shallow burial depths (< 1 km),
749
and developed along structural lineaments by hypersaline fluids presenting higher
750
temperatures and pressures than those of the host formation, which are usually
751
carbonate rocks. According to these authors, reservoir facies formed by structurally-
752
controlled hydrothermal dolomitization (HTD) are the largest hydrocarbon producers in
753
North America and have great potential in other basins. King and Goldstein (2018)
32 754
interpreted basin-derived hydrothermal system as responsible for baroque dolomite
755
precipitation with elevated Th and salinities in the Cambrian-Ordovician Arbuckle
756
Group. Hydrothermal calcite cement filling secondary porosity in hydrothermal
757
dolomites has also been described by some authors (e.g., Coveney Jr. et al. 2000,
758
Lonnee and Machel, 2006, Biehl et al., 2016, Mansurbeg et al., 2016).
759
Hydrothermal dolomitization has been reported in the Ordovician of the United
760
States, Canada, and China (e.g., Al-Aasm, 2003; Davies and Smith, 2006; Luczaj, 2006;
761
Luczaj et al., 2006; Conliffe et al., 2010; Xu et al., 2015), the Devonian and
762
Mississippian basins of Canada (e.g., Boreen and Colquhoun, 2001; Boreen and Davies,
763
2004), the Carboniferous of the United States and Spain (e.g., Gasparrini et al., 2006;
764
Hiemstra and Goldstein, 2015), the Permian of Germany (e.g., Biehl et al., 2016), the
765
Mesozoic of the Atlantic passive margins (e.g., Wierzbicki et al., 2006), the Jurassic-
766
Cretaceous of Spain, the Persian Gulf in Iraq (e.g., Mansurbeg et al., 2016), and Saudi
767
Arabia (e.g., Cantrell et al., 2004). Based on petrography, isotopic, fluid inclusions
768
characteristic, and on the operational definition previously proposed, the analyzed
769
saddle or baroque dolomites (Figs. 5, 6, 7, 8, 9, 11, 12, 13, 14) are defined as
770
hydrothermal, as also the extensive dolomitization observed in part of the associated
771
Pre-Salt rocks.
772
5.3.2.
Hydrothermal silicification
773
Silicification is a common diagenetic process, affecting a wide variety of
774
originally non-siliceous sediments (e.g., Namy, 1974; Meyers, 1977; Meyers and James,
775
1978; Hesse, 1989). Hydrothermal silicification can affect a wide variety of rock types
776
in multiple environments and geological contexts. However, few studies specifically
777
addressing the hydrothermal silicification of carbonate deposits have been published
778
(e.g., Bellanca et al., 1984; You et al., 2018). In most studies on the hydrothermal
33 779
alteration of carbonate sequences, silicification is considered a subordinate process in
780
comparison to dolomitization (e.g., Packard et al., 2001). Nevertheless, hydrothermal
781
silicification of Pre-Salt carbonates has been previously described in Campos Basin
782
(Vieira de Luca et al., 2017; Lima and De Ros, 2019), as well as of the African basins
783
(Poros et al., 2017; Teboul et al., 2017).
784
An example of the association of hydrothermal silicification and HTD
785
corresponds to the Upper Devonian (Famenian) Wabamun Group gas-condensate
786
reservoirs of the Parkland Field, Canada, which were generated by hydrothermal
787
dolomitization, silicification and dissolution (Packard et al., 2001). In those reservoirs,
788
hydrothermal silica occurs as lenses or as diffuse and discontinuous masses of
789
microquartz associated or not with dolomitization, as well as microquartz cement filling
790
fractures after saddle dolomite precipitation. The dolomitization and silicification of
791
those carbonate reservoirs occurred in relatively rapid succession and at shallow burial
792
depths. The pervasive precipitation of microcrystalline silica, macrocrystalline quartz
793
and spherulitic chalcedony observed in the study area (Figs. 5, 6, 7, 9, 11, 12, 14) was
794
defined as hydrothermal, based on the petrographic and fluid inclusions characteristics,
795
and on the operational definition previously presented.
796
5.4. Timing and source of the hydrothermal system
797
The hydrothermal processes recorded in northern Campos Basin Pre-Salt
798
lacustrine carbonates are imprinted in rocks of the sag section, and therefore cannot be
799
related to the rift magmatism. The scarcity of hydraulic breccias, the distribution of the
800
hydrothermal phases in the host carbonates, and the high homogenization temperature
801
(Th) of FIs indicate a hydrothermal system was active during effective burial, and not
802
related to the Aptian magmatism either. This indicates that the hydrothermal fluids
34 803
responsible for dissolution and for precipitation of mineral phases filling fractures that
804
cross-cut the studied carbonates are possibly related to Late Cretaceous and/or
805
Paleogene magmatic events.
806
The timing of the hydrothermal processes affecting northern Campos Basin Pre-
807
Salt deposits can be better evaluated within the framework of the burial and thermal
808
history of the area (Fig. 15). Based on this, and on petrographic and FIs evidence, the
809
observed hydrothermal alterations can be more probably related to the Late Cretaceous
810
(Santonian/Campanian), Paleocene and/or Eocene magmatic events (Fig. 15). The burial
811
and thermal history of the area indicates that during the advent of these magmatic
812
activities, the Pre-Salt reservoirs were at 2 to 4 kilometers below the seafloor and at 76
813
to 98 °C (Fig. 15). This temperature range is substantially lower (> 5–10 °C) than those
814
recorded in aqueous FIs hosted in the hydrothermal phases, thus indicating another heat
815
source to the system besides burial also had existed.
816
The δ13C data of hydrothermal SD and MC are similar (between -1.88 and
817
+1.28‰; Fig. 13, Table 2), suggesting both mineral phases were precipitated by the
818
same hydrothermal fluids, and from the same carbon source. Except for a few samples,
819
these hydrothermal phases are within the range of δ13C values of host rock mineral
820
phases (MBCB, MBCL, MBDD, FC and CS), which vary between 0.05‰ and 2.43‰.
821
Therefore, the carbon composition of the hydrothermal fluids was near to isotope
822
equilibrium with the host rock mineral phases. The slightly wider range of δ13C values
823
observed in SD samples could be explained by their wide temperature formation range
824
indicated by the FI data (105 to 150 °C, Table 1). For instance, the carbon isotopic
825
fractionation between calcite and CO2 vary from 1.0‰ at 150 °C to 3.4‰ at 100 °C
826
(Bottinga, 1969). This implies that temperature decrease alone could explain the carbon
35 827
isotopic range observed in SD, and that the isotopic composition of hydrothermal fluids
828
was buffered by their interaction with the host rocks.
829
Besides temperature, the observed variation in carbon isotopes could also be
830
related to mixture of fluids from different sources. The lowest δ13C value observed in
831
SD (-1.88‰) could suggest an original carbon isotopic for the hydrothermal fluids
832
before fluid-rock interaction, which is lower than 0.05‰. One possibility is that the
833
fluid had a δ13C value of -1.88‰. However, it is likely that this value is a result of fluid
834
with lower δ13C value and a higher δ13C value from the host-rock (mineral phases
835
usually display δ13C > 0.05‰). This interaction occurs by dissolving partially the
836
carbonates present in the host-rock. The dissolution of carbonate minerals from host
837
rock could be due to the low pH and/or low partial pressure of CO2 of the original fluid
838
before interaction with the host rock. As dissolution increases, higher is the CO2 from
839
host rock in modified hydrothermal fluid. This process is sensitive to temperature as the
840
system involves the interaction of H2O and CO2.
841
A possible original CO2 source for the hydrothermal fluids would be the
842
serpentinization of upper mantle (Fig. 16), occurring below areas with strongly thinned
843
continental crust or even through direct mantle exhumation (e.g., Boillot et al., 1987;
844
Manatschal and Bernoulli, 1999; Whitmarsh et al., 2001; Kusznir and Karner, 2007).
845
Mantle exhumation can liberate large volumes of Si, Ca, Mg and CO2 to the interacting
846
fluids (Frost and Beard, 2007; Pinto et al., 2015; 2017). Additionally, the exothermic
847
chemical reactions related to serpentinization (Moody, 1976; Proskurowski et al., 2006)
848
may generate hydrothermal fluids with high content of hydrogen and light hydrocarbons,
849
mainly CH4 (Allen and Seyfried Jr., 2004). Mantle exhumation would have occurred in
850
deep offshore southern and southeastern Brazilian basins, including Campos, during the
851
initial Atlantic Ocean opening, as suggested by seismic and gravimetric evidence
36 852
(Unternehr et al., 2010; Gomes et al., 2011; Zalán et al., 2011; Kumar et al., 2013;
853
Peron-Pinvidic et al., 2013; Kukla et al., 2018). Despite the lowest δ13C value observed
854
in SD (-1.88‰) is significantly higher than mantle isotopic composition, which shows a
855
peak carbon isotope signature of -5‰ (Deines, 2002), we cannot rule out this hypothesis.
856
Nevertheless, the 87Sr/86Sr ratios of both the syngenetic and the diagenetic/hydrothermal
857
Pre-Salt carbonates are much higher than the values expected for the derivation of Sr
858
(and of Ca, which has similar geochemical behavior; Banner, 1995) from mantle
859
exhumation and serpentinization.
860
More probable source for these δ13C values involves the contribution of light
861
carbon derived from organic matter alteration by the hydrothermal fluids, or their
862
interaction with sediments with lighter isotopic composition, such as the rift carbonates
863
from the Coqueiros Formation (e.g., well CP-5 of Dias, 1998; Fig. 16), and/or laminite
864
and spherulite lithofacies similar to those from Santos Basin Barra Velha Formation
865
(e.g., well SB-2 of Farias et al., 2019). An additional mechanism for the lowest δ13C
866
values could include the alteration of organic matter by thermal sulfate reduction (TSR;
867
Machel, 2001). The presence of sulfides and sulfates derived presumably from the same
868
hydrothermal fluid could suggest a possible role of TSR during hydrothermal fluid-rock
869
interaction. However, the carbon isotope values are not as negative as expected for this
870
process, such as in the hydrothermal calcite from Cambrian-Ordovician Arbulckle
871
Group (King and Goldstein, 2018).
872
Calculated δ18Ow from hydrothermal mineral phases partially overlap the values
873
calculated for fluids in equilibrium with the host rock mineral phases (FC and CS),
874
suggesting that part of the oxygen isotope variations can be explained by processes
875
involving dissolution of host rock during hydrothermal flow, and/or mixing between the
876
host pore waters and the hydrothermal fluids. The exceptions are the lowest calculated
37 877
δ18Ow values (roughly between 3 and 5‰), which may be related to fluids of a different
878
source or with quite distinct temperature. If the fluids had different sources, those values
879
enriched in δ18Ow would suggest extensive water-rock interaction along the fluid
880
migration pathways. Considering the geology of the basin, the Pre-Cambrian basement
881
(mainly granitic-gneissic felsic rocks), volcanic rocks from the Cabiúnas Formation
882
(mainly basaltic) and sediments from the Atafona and Coqueiros formations could have
883
interacted with the hydrothermal fluids (Fig. 16). The definition of the main rocks that
884
interacted with the hydrothermal fluids is however difficult to define based exclusively
885
on oxygen isotopes.
886
As hydrothermal fluids are usually a result of mixing between their ‘original’
887
composition and the host-rock fluids and minerals, it is likely that such fluid-rock
888
interaction would influence the
889
However, the
890
significantly higher than those from hydrothermal mineral phases (SD, MC and SB) and
891
do not overlap them. In addition, SD and SB mineral phases display high total strontium
892
content (Supplementary material). Thus, it is likely that strontium concentration of the
893
hydrothermal fluids were relatively higher than the host rock before the dissolution of
894
the latter and that SD 87Sr/86Sr ratios (~0.711) represent the ratio of hydrothermal fluids
895
before fluid-rock interaction. The
896
radiogenic than those from Cretaceous seawater, than fluids associated to hydrothermal
897
alteration of predominantly mafic igneous rocks worldwide (~0.7073; e.g., Burke et al.,
898
1982; Allègre et al., 2010), and than most of rift volcanic and volcaniclastic rocks from
899
Campos Basin Cabiúnas Formation (~0.708; Mizusaki et al., 1992; Tedeschi, 2017).
900 901
87
87
Sr/86Sr ratios of hydrothermal phases as well.
Sr/86Sr ratios of host rock mineral phases (FC, CS and BD) are
Assuming similar
87
87
Sr/86Sr ratios of SD are substantially more
Sr/86Sr ratios for the rift Coqueiros Formation to those of
Itapema Formation from Santos Basin (0.7105 to 0.7120; Tedeschi, 2017; Pietzsch et al.,
38 902
2018), such radiogenic values are likely to be related to the dissolution of the Coqueiros
903
Formation carbonates (Fig. 16). However, as the rift Itapema Formation has
904
ratios lower than those from the sag Barra Velha Formation (Pietzsch et al., 2018),
905
another source of fluids would be needed to explain such
906
syngenetic sag carbonates. The higher
907
diagenetic sag carbonates could represent a product of leaching of the Precambrian
908
felsic basement (Tupinambá et al., 2012; Teboul et al., 2016; Pietzsch et al., 2018). The
909
question is that the granitic-gneissic basement is not a suitable source for the huge
910
amounts of Mg and Ca precipitated in the sag rocks. Nevertheless, the lower
911
ratios observed in the analyzed hydrothermal phases in relation to the syngenetic and
912
early diagenetic sag carbonates could represent mixing with original fluids derived from
913
mantle serpentinization (Fig. 16), and/or a product of interaction with mafic volcanic
914
and volcanoclastic rocks from Cabiúnas Formation (Bertani and Carozzi, 1985a, b;
915
Misuzaki et al., 1992). Such mixing would be necessary to lower the 87Sr/86Sr ratios of
916
the hydrothermal mineral phases in comparison to the host rocks.
87
87
87
Sr/86Sr
Sr enrichment in the
Sr/86Sr ratios of the syngenetic and early
87
Sr/86Sr
917
According to the burial-thermal history (Fig. 15), the estimated maximum
918
temperature for the Pre-Salt reservoirs was approximately 110 °C during the occurrence
919
of the Eocene magmatic episode in the Campos Basin. This temperature is significantly
920
lower than the mean and maximum Th values, respectively 123.4 and 152 °C, obtained
921
in the aqueous FIs. The homogenization temperatures obtained in the hydrothermal
922
phases SD, MC, MQ and SB are even higher than the maximum burial temperatures
923
interpreted for the study area, including the present reservoir temperature. Therefore, the
924
Th data from SD, MC, MQ and SB represent further evidence of the hydrothermal
925
origin of these mineral phases, regardless their relation to specific magmatic periods.
926
Hydrocarbon FIs in these late-stage mineral phases suggest an atypical oil migration
39 927
(sensu Magoon and Dow, 1994) occurred in association with the hydrothermal fluid
928
flow. The absence of coexisting vapor- and liquid-rich FIs, diagnostic of phase-
929
separation effects (e.g., Goldstein and Reynolds, 1994; Jones et al., 1996; Moore et al.,
930
2001), indicates the lack of boiling processes within the studied Pre-Salt hydrothermal
931
system.
932
6. Conclusions
933
The overall geological and geochemical framework, combined with a specific
934
mineral paragenetic assemblage (saddle dolomite, macrocrystalline quartz, calcite, Sr-
935
barite, celestine, fluorite, dickite, sphalerite, galena, other metallic sulfides, and
936
bitumen) and FIs with corresponding range of salinities and homogenization
937
temperatures allowed us to recognize that the Pre-Salt carbonate reservoirs of the
938
northern Campos Basin were percolated by hydrothermal fluids chemically
939
(compositionally) comparable to those that formed in Mississippi Valley Type and
940
similar deposits.
941
The elevated homogenization temperatures (up to 152 °C) measured in
942
macrocrystalline calcite, mega-quartz, saddle dolomite and Sr-barite indicate
943
entrapment temperatures higher than the maximum values interpreted from the burial
944
history, confirming hydrothermal conditions during the late-stage alteration of the
945
studied reservoirs. The hydrothermal system characterized in northern Campos Basin
946
Pre-Salt reservoirs presents temperatures and salinities similar to Mississippi Valley and
947
Irish hydrothermal systems.
948
Saddle dolomite and macrocrystalline calcite precipitated from hydrothermal
949
fluids show lower δ18O values than the syngenetic and diagenetic carbonates. The
950
fascicular calcite aggregates, characteristic of Pre-Salt reservoirs, experienced strong
40 951
recrystallization under the action of the hydrothermal fluids and, therefore, present δ18O
952
values intermediate between the hydrothermal carbonates and the early diagenetic
953
spherulitic and microcrystalline calcite phases. All carbonates analyzed in the Pre-Salt
954
reservoirs of the northern Campos Basin present high 87Sr/86Sr, with lower ratios for the
955
hydrothermal saddle dolomite and macrocrystalline calcite than for the syngenetic,
956
diagenetic and recrystallized carbonates.
957
The high temperatures and salinities measured from the fluid inclusions, and the
958
isotopic data obtained from saddle dolomite, Sr-barite, macrocrystalline calcite and
959
mega-quartz indicate that the fault-focused hydrothermal system affecting the northern
960
Campos Basin Pre-Salt reservoirs probably involved mixing of fluids derived from
961
several sources, such as the interaction with the granitic-gneissic basement, the rift
962
sedimentary succession, the Late Cretaceous and Paleogene magmatism, and possibly
963
the rising and exhumation of the asthenosphere. The final hydrothermal fluid is a blend
964
in which end-members are not quantifiable.
965
The fluid inclusions, isotopic and petrographic data, such as the scarcity of
966
hydraulic breccia, integrated to the burial and thermal history of the study area are
967
strong evidence that the hydrothermal alteration of northern Campos Basin Pre-Salt
968
reservoirs occurred under relatively deep burial situation of more than 2 kilometers. The
969
studied hydrothermal alterations had strong impact on the porosity, permeability, and
970
heterogeneity, contributing, together with the associated fracturing, to the production
971
performance of the Pre-Salt reservoirs.
972
Acknowledgements
973
The results and interpretations of this paper are part of the PhD research project
974
of BEML, funded by Petróleo Brasileiro S.A. - Petrobras. The authors wish to thank
41 975
Petrobras, for supporting this study and for the opportunity to publish this paper. In
976
particular, we are extremely grateful to Gustavo Garcia, who provided useful insight
977
and discussions about the burial-thermal history. We are grateful for the analytical and
978
technical assistance provided by Petrobras Research Center (CENPES), the Federal
979
University of Goiás (UFG) Regional Center for Technological Development and
980
Innovation (CRTI), the University of Brasília (UnB) Geochronology Laboratory, and
981
the University of São Paulo (USP). We would also like to thank the support of the
982
Graduate Geosciences Program of Rio Grande do Sul Federal University (UFRGS). The
983
authors wish also to thank the reviewers for their constructive suggestions that helped
984
improve the manuscript.
985
References
986
Al-Aasm, 2003. Origin and characterization of hydrothermal dolomite in the Western
987
Canada Sedimentary Basin. Journal of Geochemical Exploration 78–79, 9–15.
988
Allègre, C.J., Louvat, P., Gaillardet, J., Meynadier, L., Rad, S., Capmas, F., 2010. The
989
fundamental role of island arc weathering in the oceanic Sr isotope budget. Earth and
990
Planetary Science Letters 292 (1–2), 51–56.
991
Allen, D.G., Seyfried Jr., W.E, 2004. Serpentinization and heat generation: constraints
992
from Lost City and Rainbow hydrothermal systems. Geochimica et Cosmochimica
993
Acta 68 (6), 1347–1354.
994
Alvarenga, R.S., Iacopini, D., Kuchle, J., Scherer, C.M.S., Goldberg, K., 2016. Seismic
995
characteristics and distribution of hydrothermal vent complexes in the Cretaceous
996
offshore rift section of the Campos Basin, offshore Brazil. Marine and Petroleum
997
Geology 74, 12–25.
42 998 999
Anderson, G.M., 1991. Organic maturation and ore precipitation in Southeast Missouri. Economic Geology 86 (5), 909–926.
1000
Ando, A., 2015. Intersite discrepancy in the amplitude of marine negative δ13C
1001
excursion at the onset of early Aptian oceanic anoxic event 1a: Reconciliation
1002
through Sr isotopic screening of peculiar diagenetic overprint on the Pacific
1003
reference section (Deep Sea Drilling Project Site 463). In: Neal, C.R., Sager, W.W.,
1004
Sano, T., Erba, E. (Eds.), Environmental consequences of Ontong Java Plateau and
1005
Kerguelen Plateau volcanism. Geological Society of America Special Paper 511. The
1006
Geological Society of America, Boulder, USA, pp. 329–339.
1007
Armelenti, G., Goldberg, K., Kuchle J., De Ros, L.F., 2016. Deposition, diagenesis and
1008
reservoir potential of non-carbonate sedimentary rocks from the rift section of
1009
Campos Basin, Brazil. Petroleum Geoscience 22 (3), 223–239.
1010 1011
Banner, J.L., 1995. Application of the trace element and isotope geochemistry of strontium to studies of carbonate diagenesis. Sedimentology 42 (5), 805–824.
1012
Baumgarten, C.S., Dutra, A.J.C., Scuta, M.S., Figueiredo, M.V.L., Sequeira, M.F.P.B.,
1013
1988. Coquinas da Formação Lagoa Feia, Bacia de Campos: evolução da geologia de
1014
desenvolvimento. Boletim de Geociências da Petrobras 2 (1), 27–36.
1015
Bellanca, A., Censi, P., Di Salvo, P., Neri, R., 1984. Textural, Chemical and Isotopic
1016
Variations Induced by Hydrothermal Fluids on Mesozoic Limestones in
1017
Northwestern Sicily. Mineralium Deposita 19, 78–85.
1018
Bertani, R.T., Carozzi, A.V., 1985a. Lagoa Feia Formation (Lower Cretaceous),
1019
Campos Basin, offshore Brazil: rift valley stage lacustrine carbonate reservoirs, I.
1020
Journal of Petroleum Geology 8, 37–58.
43 1021
Bertani, R.T., Carozzi, A.V., 1985b. Lagoa Feia Formation (Lower Cretaceous) Campos
1022
Basin, offshore Brazil: rift valley type lacustrine carbonate reservoirs, II. Journal of
1023
Petroleum Geology 8, 199–220.
1024
Bertrand, R., Chagnon, A., Malo, M., Duchaine, Y., Lavoie, D., Savard, M.M., 2003.
1025
Sedimentologic, diagenetic and tectonic evolution of the Saint-Flavien gas reservoir
1026
at the structural front of the Quebec Appalachians. Bulletin of Canadian Petroleum
1027
Geology 51 (2), 126–154.
1028
Biehl, B.C., Reuning, L., Schoenherr, J., Lüders, V., Kukla, P.A., 2016. Impacts of
1029
hydrothermal dolomitization and thermochemical sulfate reduction on secondary
1030
porosity creation in deeply buried carbonates: a case study from the Lower Saxony
1031
Basin, northwest Germany. American Association of Petroleum Geologists Bulletin
1032
100 (4), 597–621.
1033
Bodin, S., Meissner, P., Janssen, N.M.M., Steuber, T., Mutterlose, J., 2015. Large
1034
igneous provinces and organic carbon burial: Controls on global temperature and
1035
continental weathering during the Early Cretaceous. Global and Planetary Change
1036
133, 238–253.
1037 1038
Bodnar, R.J., 2003. Reequilibration of fluid inclusions. Fluid inclusions: Analysis and Interpretation, pp. 213–230.
1039
Bodnar, R.J., Reynolds, T.J., Kuehn, C.A., 1985. Fluid-inclusion systematics in
1040
epithermal systems. In: Berger, B., Bethke, R. (Eds.), Geology and Geochemistry of
1041
Epithermal Systems. Society of Economic Geologists, Reviews in Economic
1042
Geology 2, pp. 73–97.
1043
Bodnar, R.J., Vityk, M.O., 1994. Interpretation of microthermometric data for H2O-
1044
NaCl fluid inclusions. In: De Vivo, B., Frezzotti, M.L. (Eds.), Fluid Inclusions in
1045
Minerals, Methods and Applications. Virginia Tech Blackburg, pp. 117–130.
44 1046
Boillot, G., Recq, M., Winterer, E.L., Meyer, A.W., Applegate, J., Baltuck, M., Bergen,
1047
J.A., Comas, M.C., Davies, T.A., Dunham, K., Evans, C.A., Girardeau, J., Goldberg,
1048
G., Haggerty, J., Jansa, L.F., Johnson, J.A., Kasahara, J., Loreau, J.P., Luna-Sierra,
1049
E., Moullade, M., Ogg, J., Sarti, M., Thurow, J.,Williamson, M., 1987. Tectonic
1050
denudation of the upper mantle along passive margins: a model based on drilling
1051
results (ODP leg 103, western Galicia margin, Spain). Tectonophysics 132, 335–342.
1052
Boreen, T., Colquhoun, K., 2001. Ladyfern, NEBC: major gas discovery in the
1053
Devonian Slave Point Formation. Canadian Society of Petroleum Geologists, Annual
1054
Convention, Abstracts, 112–115.
1055
Boreen, T., Davies, G.R., 2004. Hydrothermal dolomite and leached limestones in a
1056
TCF gas play: the Ladyfern Slave Point reservoir, NEBC. In: McAuley, R. (Ed.),
1057
Dolomites - the spectrum: mechanisms, models, reservoir development. Canadian
1058
Society of Petroleum Geologists, Seminar and Core Conference, June 13–15,
1059
Calgary, Alberta, p. 17.
1060
Bottinga, Y., 1969. Calculated fractionation factors for carbon and hydrogen isotope
1061
exchange in the system calcite-carbon dioxide-graphite-methane-hydrogen-water
1062
vapor. Geochimica et Cosmochimica Acta 33, 49–64.
1063
Burke, W.H., Denison, R.E., Hetherington, E.A., Koepnick, R.B., Nelson, H.F., Otto,
1064
J.B., 1982. Variation of seawater 87Sr/86Sr throughout Phanerozoic time. Geology 10
1065
(10), 516–519.
1066
Callot, J.P., Breesch, L., Guilhaumou, N., Roure, F., Swennen, R., Vilasi, N., 2010.
1067
Paleofluids characterization and fluid flow modelling along a regional transect in
1068
Northern United Arab Emirates (UAE). Arabian Journal of Geosciences 3, 413–437.
1069 1070
Cantrell, D.L., Swart, P.K., Hagerty, R.M., 2004. Genesis and characterization of dolomite, Arab-D reservoir, Ghawar field, Saudi Arabia. GeoArabia 9 (2), 1–26.
45 1071
Castro, J.C., 2006. Evolução dos conhecimentos sobre as coquinhas-reservatório da
1072
Formação Lagoa Feia no trend Badejo-Linguado-Pampo, Bacia de Campos. Boletim
1073
de Geociências da Petrobras 25 (2), 175–186.
1074
Conliffe, J., Azmy, K., Gleeson, S.A., Lavoie, D., 2010. Fluids associated with
1075
hydrothermal dolomitization in St. George Group, western Newfoundland, Canada.
1076
Geofluids 10, 422–437.
1077
Coplen, T.B., Brand, W.A., Gehre, M., Gröning, M., Meijer, H.A.J., Toman, B.,
1078
Verkouteren, R.M., 2006. New guidelines for δ13C measurements. Analytical
1079
Chemistry 78 (7), 2439–2441.
1080
Corbella, M., Ayora, C., Cardellach, E., Soler, A., 2006. Reactive transport modeling
1081
and hydrothermal karst genesis: the example of the Rocabruna barite deposit (Eastern
1082
Pyrenees). Chemical Geology 233, 113–125.
1083
Coveney Jr., R.M., Ragan, V.M., Brannon, J.C., 2000. Temporal benchmarks for
1084
modeling Phanerozoic flow of basinal brines and hydrocarbons in the southern
1085
Midcontinent based on radiometrically dated calcite. Geology 28 (9), 795–798.
1086
Davies, G.R., 2002. Thermobaric dolomitization: Transient pressure driven processes
1087
and the role of boiling-effervescence in fabric/reservoir development (abs.).
1088
Canadian Society of Petroleum Geologists, 75th Anniversary Convention Abstracts
1089
Vol., p. 105.
1090
Davies, G.R., 2004. Hydrothermal (thermobaric) dolomitization: rock fabrics and
1091
organic petrology. In: McAuley, R. (Ed.), Dolomites - the spectrum: mechanisms,
1092
models, reservoir development. Canadian Society of Petroleum Geologists, Seminar
1093
and Core Conference, June 13–15, Calgary, Alberta, p. 17.
46 1094
Davies, G.R., Smith, L.B.J., 2006. Structurally controlled hydrothermal dolomite
1095
reservoir facies: an overview. American Association of Petroleum Geologists
1096
Bulletin 90 (11), 1641–1690.
1097 1098
Deines, P. 2002. The carbono isotope geochemistry of mantle xenoliths. Earth-Science Reviews 58 (1-4), 247–278.
1099
Dias, J.L., 1998. Análise sedimentológica e estratigráfica do Andar Aptiano em parte da
1100
Margem Leste do Brasil e no Platô das Malvinas: considerações sobre as primeiras
1101
incursões e ingressões marinhas do Oceano Atlântico Sul Meridional. (PhD Thesis).
1102
Universidade Federal do Rio Grande do Sul, Porto Alegre, Brazil, p. 208 (in
1103
Portuguese).
1104
Dias, J.L., Oliveira, J., Vieira, J., 1988. Sedimentological and stratigraphic analysis of
1105
the Lagoa Feia Formation, rift phase of Campos Basin, offshore Brazil. Revista
1106
Brasileira de Geociências 18, 252–260.
1107
Disnar, J.R., 1996. A comparison of mineralization histories for two MVT deposits,
1108
Treves and Les Malines (Causses Basin, France), based on the geochemistry of
1109
associated organic matter. Ore Geology Reviews 11 (1-3) 133–156.
1110 1111
Epstein, S., Mayeda, T., 1953. Variation of 18O content of waters from natural sources. Geochimica et Cosmochimica Acta, 4, 213–224.
1112
Farias, F., Szatmari, P., Bahniuk, A., França, A.B., 2019. Evaporitic carbonates in the
1113
pre-salt of Santos Basin – Genesis and tectonic implications. Marine and Petroleum
1114
Geology 105, 251–272
1115
Fontana, S., Nader, F.H., Morad, S., Ceriani, A., Al-Aasm, I.S., 2010. Diagenesis of the
1116
Khuff Formation (Permian-Triassic), northern United Arab Emirates. Arabian
1117
Journal of Geosciences 3 (4), 351–368.
47 1118 1119
Frost, B.R., Beard, J.S., 2007. On silica activity and serpentinization. Journal of Petrology, 48 (7), 1351–1368.
1120
Gasparrini, M., Bechstaedt, T., Boni, M., 2006. Massive hydrothermal dolomitization in
1121
the Southwestern Cantabrian Zone (Spain) and its relation to the late Variscan
1122
evolution. Marine Petroleum Geology 23, 543–568.
1123
Gilbert, G.K., 1875. Report on the geology of portions of Nevada, Utah, California, and
1124
Arizona. United States Geographic and Geological Surveys W. 100th Meridian 3,
1125
17–187.
1126
Girard, J.P., San Miguel, G., 2017. Evidence of High Temperature Hydrothermal
1127
Regimes in the Pre-Salt Series, Kwanza Basin, Offshore Angola. American
1128
Association of Petroleum Geologists Annual Convention and Exhibition. Houston,
1129
Texas, USA, Abstracts.
1130 1131
Goldstein, R.H., 2001. Fluid inclusions in sedimentary and diagenetic systems. Lithos 55, 159–193.
1132
Goldstein, R.H., Reynolds, T.J., 1994. Fluid Inclusion Microthermometry. In: Goldstein,
1133
R.H., Reynolds, T.J. (Eds.), Systematics of Fluid Inclusions in Diagenetic Minerals.
1134
Society for Sedimentary Geology, SEPM Short Courses 31, Tulsa, OK, USA, pp.
1135
87–121.
1136
Gomes, P.O., Kilsdonk, W., Grow, T., Minken, J., Barragan, R., 2011. Tectonic
1137
evolution of the Outer High of Santos Basin, southern São Paulo Plateau, Brazil, and
1138
implications for hydrocarbon exploration. In: Gao, D. (Ed.), Tectonics and
1139
sedimentation: Implications for petroleum systems. American Association of
1140
Petroleum Geologists, USA, Memoir 100, pp. 1–14.
48 1141
Greenwood, P.F., Brocks, J.J., Grice, K., Schwark, L., Jaraula, C.M.B., Dick, J.M.,
1142
Evans, K.A., 2013. Organic geochemistry and mineralogy. I. Characterization of
1143
organic matter associated with metal deposits. Ore Geology Reviews 50, 1–27.
1144 1145 1146 1147 1148 1149
Gregg, J.M., Sibley, D.F., 1984. Epigenetic dolomitization and the origin of xenotopic dolomite texture. Journal of Sedimentary Petrology 54, 908–931. Gregg, J.M., Sibley, D.F., 1987. Classification of dolomite rock textures. Journal of Sedimentary Petrology 57 (6), 967–975. Henley, R.W., Ellis, A.J., 1983. Geothermal systems ancient and modern: a geochemical review. Earth-Science Reviews 19, 1–50.
1150
Herlinger Jr., R., Zambonato, E.E., De Ros, L.F., 2017. Influence of diagenesis on the
1151
quality of lower cretaceous Pre-Salt lacustrine carbonate reservoirs from northern
1152
Campos Basin, offshore Brazil. Journal of Sedimentary Research 87, 1285–1313.
1153
Hesse, R., 1989. Silica diagenesis: origin of inorganic and replacement cherts. Earth-
1154
Science Reviews 26, 253–284.
1155
Hiemstra, E.J., Goldstein, R.H., 2015. Repeated injection of hydrothermal fluids into
1156
downdip carbonates: a diagenetic and stratigraphic mechanism for localization of
1157
reservoir porosity, Indian Basin Field, New Mexico, USA. In: Agar, S.M., Geiger, S.
1158
(Eds.), Fundamental controls on fluid flow in carbonates: current workflows to
1159
emerging technologies. Geological Society of London, Special Publications 406, pp.
1160
141–177.
1161
Holmes, A., 1928. The Nomenclature of Petrology. Ed. Van Nostrand-Reinhold, 284 p.
1162
Ilchik R.P., Brimhall G.H., Schull H.W., 1986. Hydrothermal maturation of indigenous
1163
organic matter at the Alligator Ridge Gold Deposits, Nevada. Economic Geology 81
1164
(1), 113–130.
49 1165 1166
Inoue, A., 1995. Formation of clay minerals in hydrothermal environments. In: Velde, B. (Ed.), Origin and Mineralogy of Clays. Springer-Verlag, Berlin, pp. 268−330.
1167
Jenkyns, H.C., Paull, C.K., Cummins, D.I., Fullagar, P.D., 1995. Strontium isotope
1168
stratigraphy of Lower Cretaceous atoll carbonates in the Mid-Pacific Mountains. In:
1169
Winterer, E.L., Sager, W.W., Firth, J.V., Sinton, J.M. (Eds.), Proceedings of the
1170
Ocean Drilling Program, Scientific Results 143. Texas A & M University, Ocean
1171
Drilling Program, College Station, USA, pp. 89–97.
1172
Jones, B., Renaut, R.W., Rosen, M.R., 1996. High-temperature (>90°C) calcite
1173
precipitation at Waikite Hot Springs, North Island, New Zealand. Journal of the
1174
Geological Society 153, 481–496.
1175
Khaska, M., Le Gal La Salle, C., Lancelot, J., team, ASTER, Mohamad, A., Verdoux,
1176
P., Noret, A., Simler, R., 2013. Origin of groundwater salinity (current seawater vs.
1177
saline deep water) in a coastal karst aquifer based on Sr and Cl isotopes. Case study
1178
of the La Clape massif (southern France). Applied Geochemistry 37, 212–227.
1179
King, B.D., Goldstein, R.H., 2018. History of hydrothermal fluid flow in the
1180
midcontinent,
1181
unconformities and porosity distribution. In: Armitage, P.J., Butcher, A.R., Churchill,
1182
J.M., Csoma, A.E., Hollis, C., Lander, R.H., Omma, J.E., Worden, R.H. (Eds.),
1183
Reservoir Quality of Clastic and Carbonate Rocks: Analysis, Modelling and
1184
Prediction. Geological Society of London, Special Publications 435, pp. 283–320.
1185
Kukla, P.A., Strozyk, F., Mohriak, W.U., 2018. South Atlantic salt basins – Witnesses
1186
USA:
the
relationship
between
inverted
thermal
structure,
of complex passive margin evolution. Gondwana Research 53, 41−57.
1187
Kumar, N., Danforth, A., Nuttall, P., Helwig, J., Bird, D.E., Venkatraman, S., 2013.
1188
From oceanic crust to exhumed mantle: a 40 year (1970–2010) perspective on the
1189
nature of crust under the Santos Basin, SE Brazil. In: Mohriak, W.U., Danforth, A.,
50 1190
Post, P.J., Brown, D.E., Tari, G.C., Nemcok, M., Sinha, S.T. (Eds.), Conjugate
1191
Divergent Margins. Geological Society of London, Special Publications 369, pp.
1192
147–165.
1193
Kusznir, N.J., Karner, G.D., 2007. Continental lithospheric thinning and breakup in
1194
response to upwelling divergent mantle flow: application to the Woodlark,
1195
Newfoundland and Iberia margins. In: Karner, G.D., Manatschal, G., Pinheiro, L.M.
1196
(Eds.), Imaging, Mapping and Modelling Continental Lithosphere Extension and
1197
Breakup. Geological Society of London, Special Publications 282, pp. 389–419.
1198
Leach, D.L., 1994. Genesis of the Ozark Mississippi Valley-type metallogenic province.
1199
In: Fontboté, L., Boni, M., (Eds.), Sediment Hosted Zn-Pb Ores. Springer-Verlag, pp.
1200
104–138.
1201
Leach, D.L., Bradley, D., Lewchuk, M.T., Symons, D.T.A., De Marsily, G., Brannon, J.,
1202
2001. Mississippi Valley-type lead-zinc deposits through geological time:
1203
Implications from recent age-dating research. Mineralium Deposita 36, 711–740.
1204
Lepley, S., Piccoli, L., Chitale, V., Kelley, I., Quest, M., 2017. The Importance of
1205
Understanding Diagenesis for the Development of Pre-Salt Lacustrine Carbonates.
1206
American Association of Petroleum Geologists Annual Convention and Exhibition.
1207
Houston, Texas, USA, Abstracts.
1208 1209
Leyden, R., Asmus, H., Zembruscki, S., Bryan, G., 1976. South Atlantic diapiric structures. American Association of Petroleum Geologists Bulletin 60 (2), 196–212.
1210
Lima, B.E.M., De Ros, L.F., 2019. Deposition, diagenetic and hydrothermal processes
1211
in Aptian Pre-Salt lacustrine carbonate reservoirs of the northern Campos Basin,
1212
offshore Brazil. Sedimentary Geology 383, 55–81.
1213
Lonnee, J., Machel, H.G., 2006. Pervasive dolomitization with subsequent hydrothermal
1214
alteration in the Clarke Lake gas field, Middle Devonian Slave Point Formation,
51 1215
British Columbia, Canada. American Association of Petroleum Geologists Bulletin
1216
90, 1739–1761.
1217
Luczaj, J.A., 2006. Evidence against the Dorag (mixing-zone) model for dolomitization
1218
along the Wisconsin arch – A case for hydrothermal diagenesis. American
1219
Association of Petroleum Geologists Bulletin 90 (11), 1719–1738.
1220
Luczaj, J.A., Harrison, W.B., Williams, N.S., 2006. Fractured hydrothermal dolomite
1221
reservoirs in the Devonian Dundee Formation of the central Michigan Basin.
1222
American Association of Petroleum Geologists Bulletin 90 (11), 1787–1801.
1223 1224
Machel, H.G., 2001. Bacterial and thermochemical sulfate reduction in diagenetic settings - old and new insights. Sedimentary Geology 140, 143–175.
1225
Machel, H.G., 2004. Concepts and models of dolomitization: a critical reappraisal. In:
1226
Braithwaite, C.J.R., Rizzi, G., Darke, G. (Eds.), The geometry and petrogenesis of
1227
dolomite hydrocarbon reservoirs, vol. 235, pp. 7–63. Geological Society of London,
1228
Special Publications.
1229 1230
Machel, H.G., Lonnee, J., 2002. Hydrothermal dolomite: a product of poor definition and imagination. Sedimentary Geology 152, 163–171.
1231
Machel, H.G., Mountjoy, E.W., 1987. General constraints on extensive pervasive
1232
dolomitization - and their application to the Devonian carbonates of western Canada.
1233
Bulletin of Canadian Petroleum Geology 35 (2), 143–158.
1234
Magoon, L.B., Dow, W.G., 1994. The Petroleum System. In: Magoon, L.B, Dow, W.G.
1235
(Eds.), Chapter 1: The petroleum system - from source to trap. American Association
1236
of Petroleum Geologists, USA, Memoir 60, pp. 3–24.
1237 1238
Manatschal, G., Bernoulli, D., 1999. Architecture and tectonic evolution of nonvolcanic margins: present day Galicia and ancient Adria. Tectonics 18, 1099–1119.
52 1239
Mansurbeg, H., Morad, D., Othmana, R., Morad, S., Ceriani, A., Al-aasm, I., Kolo, K.,
1240
Spirov, P., Proust, J.N., Preat, A., Koyi, H., 2016. Hydrothermal dolomitization of
1241
the Bekhme formation (Upper Cretaceous), Zagros Basin, Kurdistan Region of Iraq:
1242
Record of oil migration and degradation. Sedimentary Geology 341, 147–162.
1243 1244
Mattews, A.K., Katz, A., 1977. Oxygen isotope fractionation during the dolomitization of calcium carbonate. Geochimica et Cosmochimica Acta 41, 1431–1438.
1245
McArthur, J.M., Howarth, R.J., Shields, G.A., 2012. Chapter 7 – Strontium Isotope
1246
Stratigraphy. In: Gradstein, F.M., Ogg, J.G., Schmitz, M.D., Ogg, G.M. (Eds.), The
1247
Geologic Time Scale. Elsevier, Boston, USA, pp. 127–144.
1248
Mello, M.R., Mohriak, W.U., Koutsoukos, E.A.M., Bacoccoli, G., 1994. Selected
1249
petroleum systems in Brazil. In: Magoon, L.B., Dow, W.G. (Eds.), The petroleum
1250
system - from source to trap. American Association of Petroleum Geologists, USA,
1251
Memoir 60, pp. 499–512.
1252 1253 1254
Meyers, W.J., 1977. Chertification in the Mississippian Lake Valley Formation, Sacramento Mountains, New Mexico. Sedimentology 24 (1), 75–105. Meyers, W.J., James, A.T., 1978. Stable isotopes of cherts and carbonate cements in the
1255
Lake
Valley
Formation
1256
Sedimentology 25, 105–124.
(Mississipian),
Sacramento
Mts.,
New
Mexico.
1257
Mizusaki, A.M.P., Petrini, R., Bellieni, P., Comin-Chiaramonti, P., Dias, J.L., De Min,
1258
A., Piccirillo, E.M., 1992. Basalt magmatism along the passive continental margin of
1259
SE Brazil (Campos Basin). Contributions to Mineralogy and Petrology 111 (2), 143–
1260
160.
1261
Moody, J.B., 1976. Serpentinization: a review. Lithos 9 (2), 125–138.
53 1262
Moore, J.N., Norman, D.I., Kennedy, B.M., 2001. Fluid inclusion gas compositions
1263
from an active magmatic–hydrothermal system: a case study of The Geysers
1264
geothermal field, USA. Chemical Geology 173, 3–30.
1265 1266
Morey, G.W., Niggli, P., 1913. The hydrothermal formation of silicates, a review. Journal of the American Chemical Society 35, 1086–1130.
1267
Muniz, M.C., Bosence, D.W.J., 2015. Pre-Salt microbialites from the Campos Basin
1268
(offshore Brazil): image log facies, facies model and cyclicity in lacustrine
1269
carbonates. In: Bosence, D.W.J., Gibbons, K.A., LeHeron, D.P., Morgan, W.A.,
1270
Pritchard, T., Vining, B.A. (Eds.), Microbial carbonates in space and time:
1271
implications for global exploration and production. Geological Society, London,
1272
Special Publications 418, pp. 221–242.
1273 1274
Namy, J.N., 1974. Early diagenetic cherts in the Marble Falls Group (Pennsylvanian) of central Texas. Journal of Sedimentary Petrology 44, 1262–1268.
1275
Neilson, J.E., Oxtoby, N.H., 2008. The relationship between petroleum, exotic cements
1276
and reservoir quality in carbonates - a review. Marine Petroleum Geology 25, 778–
1277
790.
1278 1279
O’Neil, J.R., Clayton, R.N., Mayeda, T.K., 1969. Oxygen isotope fractionation in divalent metal carbonates. Journal of Chemical Physics 51, 5547–5558.
1280
Ostendorf, J., Henjes-Kunst, F., Mondillo, N., Boni, M., Schneider, J., Gutzmer, J.,
1281
2015. Formation of Mississippi Valley–type deposits linked to hydrocarbon
1282
generation in extensional tectonic settings: Evidence from the Jabali Zn-Pb-(Ag)
1283
deposit (Yemen). Geology 43 (12), 1055–1058.
1284
Packard, J.J., Al-Aasm, I.S., Samson, I., Berger, Z., Davies, J., 2001. A Devonian
1285
hydrothermal chert reservoir: the 225 bcf Parkland field, British Columbia, Canada.
1286
American Association of Petroleum Geologists Bulletin 85 (1), 51–84.
54 1287
Paradis, S., Hannigan, P., Dewing, K., 2007. Mississippi Valley-type lead-zinc deposits.
1288
In: Goodfellow, W.D. (Ed.), Mineral Deposits of Canada: A Synthesis of Major
1289
Deposit-Types, District Metallogeny, the Evolution of Geological Provinces, and
1290
Exploration Methods. Geological Association of Canada, Mineral Deposits Division,
1291
Special Publication 5, pp. 185–203.
1292
Parnell, J., Baron, M., Mann, P., Carey, P., 2003. Oil migration and bitumen formation
1293
in a hydrothermal system, Cuba. Journal of Geochemical Exploration 78-79, 409–
1294
415.
1295
Peron-Pinvidic, G., Manatschal, G., Osmundsen, P.T., 2013. Structural comparison of
1296
archetypal Atlantic rifted margins: A review of observations and concepts. Marine
1297
and Petroleum Geology 43, 21–47.
1298
Pietzsch, R., Oliveira, D.M., Tedeschi, L.R., Queiroz Neto, J.V., Figueiredo, M.F.,
1299
Vazquez, J.C., Souza, R.S., 2018. Palaeohydrology of the Lower Cretaceous pre-salt
1300
lacustrine system, from rift to post-rift phase, Santos Basin, Brazil. Palaeogeography,
1301
Palaeoclimatology, Palaecology 507, 60–80.
1302
Pinto, V.H.G., Manatschal, G., Karpoff, A.M., Viana, A., 2015. Tracing mantle-reacted
1303
fluids in magma-poor rifted margins: The example of Alpine Tethyan rifted margins.
1304
Geochemistry Geophysics Geosystems 16, 3271–3308.
1305
Pinto, V.H.G., Manatschal, G., Karpoff, A.M., Ulrich, M., Viana, A.R., 2017. Seawater
1306
storage and element transfer associated with mantle serpentinization in magma-poor
1307
rifted margins: a quantitative approach. Earth and Planetary Science Letters 459,
1308
227–237.
1309
Poros, Z., Jagniecki, E., Luczaj, J., Kenter, J., Gal, B., Correa, T.S., Ferreira, E.,
1310
McFadden, K.A., Elifritz, A., Heumann, M., Johnston, M., Matt, V., 2017. Origin of
1311
silica in Pre-Salt carbonates, Kwanza Basin, Angola. AAPG Annual Convention and
55 1312
Exhibition, Houston, Texas, April 2–5. USA: AAPG Datapages/Search and
1313
Discovery, article #90291.
1314
Proskurowski, G., Lilley, M.D., Kelley, D.S., 2006. Low temperature volatile
1315
production at the Lost City Hydrothermal Field, evidence from a hydrogen stable
1316
isotope geothermometer. Chemical Geology, 229 (4), 331−343.
1317
Reed, M.H., 1997. Hydrothermal alteration and its relationship to ore fluid composition
1318
In: Barnes, H.L. (Ed.), Geochemistry of Hydrothermal Ore Deposits, 3rd edition.
1319
John Wiley and Sons, pp. 303–365.
1320 1321
Rodrigues, R., 2005. Chemostratigraphy. In: Koutsoukos, E.A.M. (Ed.), Applied Stratigraphy. Springer, Dordrecht, Netherlands, pp. 165–178.
1322
Rodriguez, C.R., Jackson, C.A.L., Rotevatn, A., Bell, R.E., Francis, M., 2018. Dual
1323
tectonic-climatic controls on salt giant deposition in the Santos Basin, offshore Brazil.
1324
Geosphere 14, 215–242.
1325 1326
Roedder, E., 1984. Fluid inclusions. Mineralogical Society of America, Reviews in Mineralogy 12, 646 p.
1327
Roedder, E., Bodnar, R.J., 1997. Fluid inclusion studies of hydrothermal ore deposits.
1328
In: Barnes, H.L. (Ed.), Geochemistry of Hydrothermal Ore Deposits. Wiley, New
1329
York, pp. 657–697.
1330
Sabato Ceraldi, T., Green, D., 2016. Evolution of the South Atlantic lacustrine deposits
1331
in response to Early Cretaceous rifting, subsidence and lake hydrology. In: Sabato
1332
Ceraldi, R.A., Hodgkinson, T., Backe, G. (Eds.), Petroleum Geoscience of the West
1333
Africa Margin. Geological Society, London, Special Publication 438, pp. 77–98.
1334
Saller, A., Rushton, S., Buambua, L., Inman, K., Mcneil, R., Dickson, J.A.D., 2016.
1335
Pre-Salt stratigraphy and depositional systems in the Kwanza Basin, offshore Angola.
1336
American Association of Petroleum Geologists Bulletin 100, 1135–1164.
56 1337
Schoenherr, J., Littke, R., Urai, J.L., Kukla, P.A., Rawahi, Z., 2007. Polyphase thermal
1338
evolution in the Infra-Cambrian Ara Group (South Oman Salt Basin) as deduced by
1339
maturity of solid reservoir bitumen. Organic Geochemistry 38 (8), 1293–1318.
1340 1341
Sharp, Z., 2007. Principles of Stable Isotope Geochemistry. Pearson Prentice Hall, Upper Saddle River, New Jersey, 344 p.
1342
Sibson, R.H., 1990. Faulting and fluid flow. In: Nesbitt, B.E. (Ed.), Fluids in
1343
tectonically active regimes of the continental crust. Mineralogical Association of
1344
Canada, Short Course Handbook 18, chapter 4, pp. 93–132.
1345 1346
Simoneit, B.R.T., 1990. Petroleum generation, an easy and widespread process in hydrothermal systems: an overview. Applied Geochemistry 5 (1-2), 3–15.
1347
Simoneit, B.R.T., 2018. Hydrothermal Petroleum. In: Wilkes, H. (Ed.), Hydrocarbons,
1348
Oils and Lipids: Diversity, Origin, Chemistry and Fate. Handbook of Hydrocarbon
1349
and Lipid Microbiology. Springer, p. 800.
1350
Skinner, B.J., 1997. Hydrothermal mineral deposits: what we do and don't know. In:
1351
Barnes, H.L. (Ed.), Geochemistry of hydrothermal ore deposits, 3rd edition. New
1352
York: Wiley and Sons, pp. 1–29.
1353
Smith, L.B.J., 2006. Origin and reservoir characteristics of Upper Ordovician Trenton–
1354
Black River hydrothermal dolomite reservoirs in New York. American Association
1355
of Petroleum Geologists Bulletin 90 (11), 1691–1718.
1356 1357
Stearns, N.D., Stearns, H.T., Waring, G.A., 1935. Thermal springs in the United States. United States Geological Survey, Water Supply Paper 679-B, 59-191.
1358
Steele-MacInnis, M., Bodnar, R.J., Naden, J., 2011. Numerical model to determine the
1359
composition of H2O-NaCl-CaCl2 fluid inclusions based on microthermometric and
1360
microanalytical data. Geochimica et Cosmochimica Acta 75, 21–40.
57 1361
Suchý, V., Dobes, P., Sýkorová, I., Machovic, V., Stejskal, M., Kroufek, J., Chudoba, J.,
1362
Matejovský, L., Havelcová, M., Maysová, P., 2010. Oil-bearing inclusions in vein
1363
quartz and calcite and, bitumens in veins: Testament to multiple phases of
1364
hydrocarbon migration in the Barrandian basin (lower Palaeozoic), Czech Republic.
1365
Marine and Petroleum Geology 27 (1), 285–297.
1366
Teboul, P.A., Durlet, C., Gaucher, E.C., Virgone, A., Girard, J.P., Curie, J., Lopez, B.,
1367
Camoin, G.F., 2016. Origins of elements building travertine and tufa: New
1368
perspectives provided by isotopic and geochemical tracers. Sedimentary Geology
1369
334, 97–114.
1370
Teboul, P.A., Durlet, C., Girard, J.P., Dubois, L., Miguel, G.S., Virgone, A., Gaucher,
1371
E.C., Camoin, G., 2019. Diversity and origin of quartz cements in continental
1372
carbonates: example from the Lower Cretaceous rift deposits of the South Atlantic
1373
margin. Applied Geochemistry 100, 22–41.
1374
Teboul, P.A., Kluska, J.M., Marty, N.C.M., Debure, M., Durlet, C., Virgone, A.,
1375
Gaucher, E.C., 2017. Volcanic rock alterations of the Kwanza Basin, offshore
1376
Angola - Insights from an integrated petrological, geochemical and numerical
1377
approach. Marine and Petroleum Geology 80, 394–411.
1378
Tedeschi, L.R., 2017. Lower Cretaceous climate records and the correlation between
1379
marine and lacustrine settings (Europe and South America). (PhD Thesis). University
1380
of Oxford, Oxford, UK, p. 343.
1381
Thompson, D.L., Stilwell, J., Hall, M., 2015. Lacustrine carbonate reservoirs from Early
1382
Cretaceous rift lakes of Western Gondwana: Pre-Salt coquinas of Brazil and West
1383
Africa. Gondwana Research 28, 26–51.
58 1384
Tosca, N.J., Wright, V.P., 2014. The formation and diagenesis of Mg-clay minerals in
1385
lacustrine carbonate reservoirs. AAPG Annual Convention and Exhibition, Houston,
1386
Texas, April 6-9. USA: AAPG Datapages/Search and Discovery, article #51002.
1387
Tupinambá, M., Heilbron, M., Valeriano, C., Júnior, R.P., Dios, F.B., Machado, N.,
1388
Silva, L.G.D.E., Almeida, J.C.H., 2012. Juvenile contribution of the Neoproterozoic
1389
Rio Negro Magmatic Arc (Ribeira Belt, Brazil): Implications for Western Gondwana
1390
amalgamation. Gondwana Research 21 (2–3), 422–438.
1391
Unternehr, P., Péron-Pinvidic, G., Manatschal, G., Sutra, E., 2010. Hyper-extended
1392
crust in the South Atlantic: in search of a model. Petroleum Geoscience 16, 207–215.
1393
Utada, M, 1980. Hydrothermal alterations related to igneous activity in Cretaceous and
1394
Neogene formations of Japan. Mining Geology, Special Issue 8, 67–83.
1395
Ventura, G.T., Simoneit, B.R.T., Nelson, R.K., Reddy, C.M., 2012. The composition,
1396
origin and fate of complex mixtures in the maltene fractions of hydrothermal
1397
petroleum assessed by comprehensive two-dimensional gas chromatography.
1398
Organic Geochemistry 45, 48–65.
1399
Vieira de Luca, P.H., Matias, H., Carballo, J., Sineva, D., Pimentel, G.A., Tritlla, J.,
1400
Esteban, M., Loma, R., Alonso, J.L.A., Jiménez, R.P., Pontet, M., Martinez, P.B.,
1401
Vega, V., 2017. Breaking barriers and paradigms in presalt exploration: The Pão de
1402
Açúcar discovery (offshore Brazil). In: Merrill, R.K., Sternbach C.A. (Eds.), Giant
1403
fields of the decade 2000–2010. American Association of Petroleum Geologists,
1404
USA, Memoir 113, pp. 177–194.
1405 1406 1407 1408
White, D.E., 1957. Thermal waters of volcanic origin. Geological Society of America Bulletin 68, 1637–1658. Whitmarsh, R.B., Manatschal, G., Minshull, T.A., 2001. Evolution of magma-poor continental margins from rifting to sea-floor spreading. Nature 413, 150–153.
59 1409
Wierzbicki, R., Dravis, J.J., Al-Aasm, I., Harland, N., 2006. Burial dolomitization and
1410
dissolution of Upper Jurassic Abenaki platform carbonates, Deep Panuke reservoir,
1411
Nova Scotia, Canada. American Association of Petroleum Geologists Bulletin 90,
1412
1843–1861.
1413 1414
Wilkinson, J.J., 2001. Fluid inclusions in hydrothermal ore deposits. Lithos 55, 229−272.
1415
Wilkinson, J.J., 2010. A Review of Fluid Inclusion Constraints on Mineralization in the
1416
Irish Ore Field and Implications for the Genesis of Sediment-Hosted Zn-Pb Deposits.
1417
Economic Geology 105, 417–442.
1418
Wilson, M.I.J., Evans, M.J., Oxtoby, N.H., Nas, D.S., Donnelly, T., Thirwall, M., 2007.
1419
Reservoir quality, textural evolution, and origin of fault-associated dolomites.
1420
American Association of Petroleum Geologists Bulletin 91, 1247–1272.
1421
Wilson, N.S.F., Zentilli, M., 1999. The role of organic matter in the genesis of the El
1422
Soldado volcanic-hosted manto-type Cu deposit, Chile. Economic Geology 94 (7),
1423
1115–1136.
1424 1425 1426 1427
Winter, W.R., Jahnert, R.J., França, A.B., 2007. Bacia de Campos. Boletim de Geociências da Petrobras 15, 511–529. Wright, V.P., 2011. Reservoir architectures in non-marine carbonates. AAPG Annual Convention and Exhibition, Proceedings, Houston, p. 75–89.
1428
Wright, V.P., 2012. Lacustrine carbonates in rift settings: the interaction of volcanic and
1429
microbial processes on carbonate deposition. In: Garland, J., Neilson, J.E., Laubach,
1430
S., Whidden, K.J. (Eds.), Advances in carbonate exploration and reservoir analysis.
1431
Geological Society of London, Special Publications 370, pp. 39–47.
60 1432
Wright, V.P., Barnett, A.J., 2014. Cyclicity and Carbonate-Silicate Gel Interactions in
1433
Cretaceous Alkaline Lakes. AAPG Annual Convention and Exhibition, Houston,
1434
Texas. Search and Discovery Article #51011.
1435
Wright, V.P., Barnett, A.J., 2015. An abiotic model for the development of textures in
1436
some South Atlantic Early Cretaceous lacustrine carbonates. In: Grotzinger, J.P.,
1437
James, N. (Eds.), Microbial carbonates in space and time: implications for global
1438
exploration and production. Geological Society of London, Special Publications 418,
1439
pp. 209–219.
1440
Wright, V.P., Tosca, N.J., 2016. Geochemical Model for the Formation of the Pre-Salt
1441
Reservoirs, Santos Basin, Brazil: Implications for Understanding Reservoir
1442
Distribution. American Association of Petroleum Geologists Annual Convention and
1443
Exhibition. Calgary, Alberta, Canada.
1444
Xu, K., Yu, B., Gong, H., Ruan, Z., Pan, Y., Ren, Y., 2015. Carbonate reservoirs
1445
modified by magmatic intrusions in the Bachu area, Tarim Basin, NW China.
1446
Geoscience Frontiers 6, 779–790.
1447
Yamamoto, K., Ishibashi, M., Takayanagi, H., Asahara, Y., Sato, T., Nishi, H., Iryu, Y.,
1448
2013. Early Aptian paleoenvironmental evolution of the Bab Basin at the southern
1449
Neo-Tethys margin: Response to global carbon-cycle perturbations across Ocean
1450
Anoxic Event 1a. Geochemistry, Geophysics, Geosystems 14 (4), 1104–1130.
1451
You, D., Han, J., Hu, W., Qian, Y., Chen, Q., Xi, B., Ma, H., 2018. Characteristics and
1452
formation mechanisms of silicified carbonate reservoirs in well SN4 of the Tarim
1453
Basin. Energy Exploration & Exploitation 36 (4), 820–849.
1454
Zalán, P.V., Severino, M.C.G., Rigoti, C.A., Magnavita, L.P., Bach, J.A., 2011. An
1455
Entirely New 3D-View of the Crustal and Mantle Structure of a South Atlantic
61 1456
Passive Margin - Santos, Campos and Espírito Santo Basins, Brazil. AAPG Annual
1457
Convention and Exhibition, Proceedings, Houston, 12p.
1458
Zhu, D., Meng, Q., Jin, Z., Liu, Q., Hu, W., 2015. Formation mechanism of deep
1459
Cambrian dolomite reservoirs in the Tarim basin, northwestern China. Marine and
1460
Petroleum Geology 59, 232–144.
1461 1462
62 1463
Figures and Tables
1464 1465
Figure 1. Location map of the study area in the northern Campos Basin, offshore
1466
southeast Brazil (modified from Dias et al., 1988; Lima and De Ros, 2019). The straight
1467
dashed lines define the limits of the Campos Basin with the Espírito Santo Basin to the
1468
North (Vitória High) and with the Santos Basin to the South (Cabo Frio High). The
1469
detail indicates the location of the four Petrobras wells used in the study.
1470 1471
Figure 2. Summarized Stratigraphic chart of the Lagoa Feia Group in Campos Basin,
1472
offshore Brazil (modified from Winter et al., 2007; Lima and De Ros, 2019). Wavy
1473
lines represent unconformities.
1474 1475
Figure 3. Detailed stratigraphic logs of the Coqueiros (rift section) and Macabu (sag
1476
phase) formations at Campos Basin, offshore Brazil, in three of the four wells used in
1477
the study. The stratigraphic datum is the top of the sag reservoir (base of evaporites
1478
from Ariri Formation, which is not shown). Dolomitization and silicification intensities
1479
are indicated by the size of the bars increasing to the left. GR: gamma ray curve; δ13C
1480
and δ18O: samples analyzed for oxygen and carbon isotopes; 87Sr/86Sr: samples analyzed
1481
for strontium isotopes; FI: samples analyzed for fluid inclusions. Red wavy lines
1482
represent unconformities.
1483 1484
Figure 4. Photomicrographs highlighting main features in rift and sag lithologic types,
1485
Campos Basin, Brazil. A) Rudstone of fragmented, partially dissolved bivalve bioclasts,
1486
with well-preserved interparticle porosity (blue) (plane-polarized light with uncrossed
1487
polarizers; //P). B) Stevensitic arenite with stevensite ooids and peloids strongly
63 1488
dissolved and silicified (white) (//P). C) Fascicular calcite crusts (CFC) intercalated
1489
with granular deposits replaced by microcrystalline dolomite (MD) (crossed polarizers;
1490
XP). D) Calcite spherulites (CS) in Mg-clay matrix replaced by microcrystalline
1491
dolomite (MD) (XP). E) Planar to wavy laminite with millimetric levels defined by
1492
predominance of clay or microcrystalline dolomite (//P). F) Intraclastic grainstone with
1493
interparticle porosity partially cemented by calcite and dolomite (//P).
1494 1495
Figure 5. Paragenetic sequences interpreted for the sag fascicular calcite crusts from
1496
northern Campos Basin, Brazil. The paragenetic sequence is divided into syngenetic
1497
(Syn), eodiagenetic, mesodiagenetic and hydrothermal phases (modified from Lima and
1498
De Ros, 2019). The thickness of the lines corresponds to the intensity or abundance of
1499
the processes and products. The red lines represent porosity loss, and the blue lines
1500
porosity increase. Tha and Tho show the mean homogenization temperatures of primary
1501
aqueous and oil fluid inclusions, respectively (minimum and maximum limits in
1502
brackets).
1503 1504
Figure 6. Schematic representation of the main syngenetic, diagenetic and hydrothermal
1505
processes of formation and alteration of the studied Aptian Pre-Salt deposits, Campos
1506
Basin, offshore Brazil (modified from Lima and De Ros, 2019). (A) Syngenetic crust of
1507
coalesced fascicular calcite aggregates covering syngenetic Mg-clay matrix partially
1508
replaced and displaced by calcite spherulites. B) Microcrystalline and blocky dolomite
1509
partially replacing the Mg-clay matrix, calcite spherulites, and fascicular aggregates.
1510
Dolomite rhombohedra partially filling inter-aggregate pores. Dissolution of the
1511
syngenetic Mg-clay matrix and part of the dolomitized calcite aggregates. C)
1512
Microcrystalline quartz and chalcedony spherulites partially to fully replacing
64 1513
preexisting constituents, and filling primary and secondary porosity. D) Silicification of
1514
some dolomitized fascicular and spherulitic aggregates. Svanbergite as pseudocubic
1515
crystals disseminated in microcrystalline silica or within chalcedony spherulites.
1516
Recrystallization of fascicular calcite and initiation of fracturing process. E) Intense
1517
silicification, fracturing and dissolution. Saddle dolomite partially filling fracture and
1518
vugular pores. F) Macrocrystalline calcite, radiated Sr-barite and celestine, fibro-
1519
radiated zeolite, euhedral to subhedral pyrite, chalcopyrite, galena and sphalerite, and
1520
massive bitumen filling the secondary pores. The numbering is in accordance with the
1521
paragenetic sequence in Figure 5.
1522 1523
Figure 7. Photomicrographs showing aspects of sag lithologic types affected by
1524
dolomitization and silicification. A) Incipiently-dissolved saddle dolomite (SD) filling
1525
vugular pore (XP). B) Macrocrystalline calcite (MC) and prismatic, elongated and
1526
oriented celestine (Clt) filling fractures in brecciated dolostone (Dol) (XP). C)
1527
Cathodoluminescence (CL) image of zoned saddle dolomite crystals (SD) partially
1528
dissolved and replaced by microcrystalline silica (MS). D) Zoned saddle dolomite
1529
intensely silicified (SSD), macrocrystalline quartz (MQ), and spherulitic chalcedony
1530
(ChS) completely filling vugular pore (XP). E) Drusiform quartz (DQ) and spherulitic
1531
chalcedony (ChS) cementing vugular porosity in partially dolomitized and silicified
1532
protolith (XP). F) Radial prismatic quartz (PQ), prismatic Sr-barite (SB),
1533
macrocrystalline calcite (MC) and euhedral pyrite (blue arrows) filling vugular porosity
1534
(XP).
1535 1536
Figure 8. Photomicrographs from petrographic features of some hydrothermal mineral
1537
phases identified in this study. A) Prismatic Sr-barite (SB), saddle dolomite (SD) and
65 1538
macrocrystalline calcite (MC) (stained red) filling vugular pore (XP). B) CL image of
1539
zoned macrocrystalline calcite (MC) and saddle dolomite (SD) filling vugular porosity.
1540
C) Elongated prismatic Sr-barite (SB) and euhedral macrocrystalline calcite (MC)
1541
partially filling vugular porosity (backscattered electrons image; BSE). D) Sr-barite
1542
(SB) filling vugular porosity in dolostone (Dol) (BSE). E) Fibro-radiated zeolite (Ze)
1543
and euhedral macrocrystalline calcite (MC) filling partially vugular pore in chert
1544
composed by microcrystalline silica (MS) (BSE). F) Blocky dolomite (BD),
1545
macrocrystalline calcite (MC) and bitumen (Bi) cementing vugular porosity in rock with
1546
totally recrystallized and partially dolomitized calcite spherulites (CS) (//P).
1547 1548
Figure 9. Photomicrographs showing examples of fluid inclusions (FIs) in double-
1549
polished sections (120 µm) showing primary aqueous and petroleum fluid inclusions
1550
(blue and red arrows, respectively) from Campos Basin Pre-Salt, Brazil. Liquid-rich
1551
aqueous inclusions hosted in: A) Macrocrystalline calcite, B) Quartz, and C) Sr-barite
1552
(//P). D) Photomicrograph shows a saddle dolomite growth zones bearing aligned
1553
primary petroleum inclusions (//P). E) Photomicrograph in (//P) and F) ultraviolet
1554
fluorescence image of saddle dolomite hosting primary aqueous and pseudo-secondary
1555
petroleum fluid inclusions.
1556 1557
Figure 10. Homogenization temperature (°C) histograms of aqueous and petroleum
1558
(hatched) fluid inclusions hosted in syngenetic and diagenetic constituents from Pre-Salt
1559
sag and rift intervals, Campos Basin, offshore Brazil. (A) Primary fluid inclusions
1560
hosted in all analyzed constituents. (B) Primary fluid inclusions hosted in recrystallized
1561
calcite bioclasts. (C) Primary fluid inclusions hosted in recrystallized fascicular calcite.
1562
(D) Primary fluid inclusions hosted in recrystallized calcite spherulites. (E) Secondary
66 1563
fluid inclusions hosted in recrystallized calcite spherulites. The average formation
1564
temperature currently measured in the depth interval where the samples for fluid
1565
inclusions were obtained is 113.7 °C.
1566 1567
Figure 11. Salinities estimated for minerals phases from Pre-Salt rift and sag intervals,
1568
Campos Basin, Brazil. Histograms of the salinities (wt. % of equivalent in NaCl),
1569
computed from NaCl-H2O system, for primary and pseudo-secondary (A) and
1570
secondary (B) fluid inclusions hosted in saddle dolomite (orange), Sr-barite (black), and
1571
macrocrystalline quartz (brown) and calcite (red) hydrothermal phases. Note that the
1572
highest salinity values were obtained in the secondary fluid inclusions. The formation
1573
water salinity sampled in the rift reservoir is 27.2 wt. % NaCl eq. (*) Salinities of
1574
aqueous fluid inclusions hosted in recrystallized fascicular calcite calculated in wt. %
1575
NaCl-CaCl2 eq.
1576 1577
Figure 12. Homogenization temperature (°C) histograms of aqueous and petroleum
1578
(hatched) fluid inclusions hosted in hydrothermal mineral phases from Pre-Salt sag and
1579
rift intervals, Campos Basin, Brazil. In the left, primary and pseudo-secondary and in
1580
the right side secondary fluid inclusions obtained in all hydrothermal phases (A and B),
1581
Sr-barite (black) (C and D), macrocrystalline quartz (brown) (E and F), macrocrystalline
1582
calcite (red) (G and H), and saddle dolomite (orange) (I). Note that the homogenization
1583
temperatures obtained from oil fluid inclusions are consistently lower than those of the
1584
aqueous fluid inclusions in all mineral phases.
1585 1586
Figure 13. Carbonate isotope composition, Campos Basin Pre-Salt, offshore Brazil.
1587
Cross-plots of: A) and B) δ18O values versus δ13C values (both in ‰ VPDB) by
67 1588
lithofacies and wells. Note the gray polygons delimiting the rift (r) and sag (s) data from
1589
Campos (Dias, 1998; Rodrigues, 2005; Muniz and Bosence, 2015), Santos (Farias et al.,
1590
2019), and Kwanza Basins (Saller et al., 2016; Sabato Ceraldi and Green, 2016). C) and
1591
D) 87Sr/86Sr ratio versus δ18O values (in ‰ VPDB). E) 87Sr/86Sr ratio versus δ13C values
1592
(in ‰ VPDB). F)
1593
rocks compared to Early Cretaceous and Phanerozoic seawater (Jenkyns et al., 1995;
1594
McArthur et al., 2012; Yamamoto et al., 2013; Bodin et al., 2015; Ando, 2015) and Pre-
1595
Salt carbonate (Dias, 1998; Tedeschi, 2017; Pietzsch et al., 2018; Farias et al., 2019).
1596
Samples of microcrystalline and blocky calcite in recrystallized bioclasts (dark green
1597
triangles), recrystallized fascicular calcite (purple squares), microcrystalline and blocky
1598
calcite in laminite (light green diamonds), recrystallized calcite spherulites (blue
1599
squares), microcrystalline and blocky dolomite (pink circles), and hydrothermal
1600
carbonates filling secondary porosity and fractures (saddle dolomite - orange circles;
1601
macrocrystalline calcite - red rhombs). Micromill and bulk samples are indicated with
1602
and without external lines, respectively. Note that the δ18O values are closer to zero for
1603
primary, syngenetic and eodiagenetic constituents, more negative for the hydrothermal
1604
phases, and intermediate for the recrystallized fascicular calcite. In the 87Sr/86Sr vs. δ18O
1605
and
1606
syngenetic and eo/mesodiagenetic constituents (blue line) and another for the
1607
hydrothermal phases (red line). In the hydrothermal paragenesis, the more radiogenic
1608
values (higher
1609
well as the correlation factors (R) are indicated.
87
Sr/86Sr values obtained in the hydrothermal phases and sag host
87
Sr/86Sr vs. δ13C cross-plots there are two well-defined trends, one for the
87
Sr/86Sr ratio) present more negative values of δ18O. The equations as
1610 1611
Figure 14. Homogenization temperature (°C) versus salinity (wt. % NaCl eq.) cross plot
1612
for aqueous fluid inclusions from Pre-Salt sag and rift samples, Campos Basin, offshore
68 1613
Brazil. Primary/pseudo-secondary (upper part) and secondary (lower part) fluid
1614
inclusions hosted in hydrothermal phases Sr-barite (black), macrocrystalline calcite
1615
(red) and quartz (brown), and saddle dolomite (orange). Salinity in primary/pseudo-
1616
secondary fluid inclusions varies in the range of 12 to 26 wt. % eq. NaCl. However,
1617
97% of the data obtained from aqueous fluid inclusions show salinities higher than 15
1618
wt. % eq. NaCl. (*) Salinity of aqueous fluid inclusions hosted in recrystallized
1619
fascicular calcite calculated in wt. % NaCl-CaCl2 eq.
1620 1621
Figure 15. Burial-thermal history diagram for northern Campos Basin Pre-Salt
1622
lacustrine reservoirs, highlighting the burial depths (black line) of the sag reservoir (S)
1623
and of the 100 °C isotherm (dashed red line). The surface temperatures estimated for rift
1624
and sag deposits are 26.9 and 27.8 °C, respectively. Color scale and red values within
1625
the white rectangles represent the temperature variation. The blue arrows in the lower
1626
right corner indicate the depths of the samples analyzed for fluid inclusions. In the left,
1627
a histogram of homogenization temperatures (Th) of primary/pseudo-secondary aqueous
1628
fluid inclusions hosted in fascicular calcite, saddle dolomite, macrocrystalline calcite,
1629
mega-quartz, and Sr-barite, indicating precipitation temperatures significantly higher
1630
than the burial temperatures. Note that subsidence was essentially continuous since the
1631
deposition of the Pre-Salt sag sequence. Magmatic activity (M) occurring in the Campos
1632
Basin during Cretaceous and Paleogene (Winter et al., 2007) is shown by purple
1633
rectangles and black dashed lines. The dashed green line indicates the peak of
1634
conventional oil generation and migration (OGM) in the Miocene of the Campos Basin,
1635
according to Mello et al. (1994).
1636
69 1637
Figure 16. Schematic representation of the deep-burial hydrothermal system affecting
1638
the northern Campos Basin Pre-Salt reservoirs, offshore Brazil. The fault-focused
1639
hydrothermal system probably involved mixing of fluids derived from several sources:
1640
(B) Pre-Cambrian basement (mainly granitic-gneissic felsic rocks), (C) volcanic rocks
1641
from the Cabiúnas Formation (mainly basaltic), (R) rift deposits from the Atafona and
1642
Coqueiros formations, and (S) serpentinization of the upper mantle. Note the intrusive
1643
magmatic (mafic) rocks (M) in the Coqueiros (rift interval) and Macabu (sag section)
1644
formations; hydrothermal alteration (H) comprising extensive dolomitization,
1645
silicification, and dissolution, with the paragenesis including saddle dolomite,
1646
macrocrystalline calcite, mega-quartz, Sr-barite, celestine, fluorite, dickite, sphalerite,
1647
galena, and other metallic sulfides filling fractures and dissolution porosity; and
1648
anhydrite thick layer (A) at the top of the sag reservoir affected by the hydrothermal
1649
alteration (base of evaporites).
1650 1651
Table 1. Statistical summary of the results obtained in aqueous and petroleum
1652
inclusions from Campos Basin Pre-Salt, offshore Brazil. Fluid inclusions are hosted in
1653
the recrystallized bioclasts and syngenetic fascicular calcite, diagenetic recrystallized
1654
calcite spherulites, saddle dolomite, macrocrystalline calcite, mega-quartz, and radiated
1655
Sr-barite. Lithologic types - br: bioclastic rudstone; fcc: fascicular calcite crust; sc:
1656
stevensitic claystones with calcite spherulites; ig: intraclastic grainstone; ct: chert.
1657
Comp.: composition; FC: fluorescence color with Nikon UV-2A filter; API hc:
1658
estimated API gravity of petroleum inclusions; Th: liquid-vapor homogenization
1659
temperature of aqueous and petroleum fluid inclusions; Tm ICE: ice melting temperature
1660
of aqueous fluid inclusions; Sal (wt% NaCl eq.): equivalent salinities estimated from
1661
final ice melting temperatures in weight percent of NaCl; n: number of inclusions
70 1662
measured; pr: primary fluid inclusions; sec: secondary fluid inclusions; psec: pseudo-
1663
secondary fluid inclusions; aq: aqueous fluid inclusions; hy: hydrocarbon fluid
1664
inclusions; wt: white fluorescence; bl: blue fluorescence; yl: yellow fluorescence; pyl:
1665
pale yellow fluorescence; mx: mixed inclusions containing aqueous fluid and
1666
petroleum; N/A: could not be determined. (*) Salinities of aqueous fluid inclusions
1667
hosted in recrystallized fascicular calcite calculated in wt. % NaCl-CaCl2 eq.
1668 1669
Table 2. δ13C and δ18O results of isotopic analyses obtained in bioclasts recrystallized
1670
to microcrystalline and blocky calcite, recrystallized fascicular calcite, laminite with
1671
microcrystalline and blocky calcite, recrystallized calcite spherulites, microcrystalline
1672
and blocky dolomite replacing pre-existing constituents, and hydrothermal saddle
1673
dolomite and macrocrystalline calcite filling dissolution porosity and fractures.
1674 87
Sr/86Sr ratio obtained from recrystallized fascicular calcite and
1675
Table 3. Results of
1676
calcite spherulites, microcrystalline and blocky dolomite, and hydrothermal saddle
1677
dolomite, macrocrystalline calcite and Sr-barite filling dissolution porosity and fractures.
1678 1679
Table 4. δ18Ow in VSMOW calculated based on Th values from fluid inclusions and
1680
measured δ18Ocarb measured in carbonates from Campos Basin Pre-Salt, Brazil. The
1681
equations of O’Neil et al. (1969) and Mathews and Katz (1977) were used to calculate
1682
δ18Ow in VPDB, which was converted to δ18Ow in VSMOW by using the conversion of
1683
O’Neil et al. (1969). Details are given in item 5.2.
1684
71 1685
Supplementary Material
1686 1687
Supplementary material 1. WDS (Wavelength-Dispersive X-Ray Spectroscopy)
1688
chemical composition of carbonates from Pre-Salt reservoir in the Campos Basin, Brazil.
1689
The box-plots show the contents of (A) Na2O, (B) Al2O3, (C) SiO2, (D) MnO, (E) FeO,
1690
(F) SrO, (G) BaO and (H) Ce2O3 in mass %, obtained from fascicular (FC; purple;
1691
n=72), spherulitic (CS; blue; n=66) and macrocrystalline (MC; red; n=140) calcite, and
1692
microcrystalline (MD; light pink; n=32), blocky (BD; dark pink; n=40) and saddle
1693
dolomite (SD; orange; n=88). Thick lines and rhombs within the boxes, and circles
1694
correspond to median, mean and outliers, respectively.
1695 1696
Supplementary material 2. Scanning electron microscopy (SEM) analyses in selected
1697
hydrothermal minerals from Campos Basin Pre-Salt, Brazil. Backscattered electrons
1698
images (BSE) and As, S, Co, Cu, Zn, Ni, Fe and Pb contents (mass%) from WDS
1699
(Wavelength-Dispersive X-Ray Spectroscopy) analyses performed on sphalerite (Sp)
1700
and euhedral pyrite (Py), and Na2O, and P2O5, SO3, CaO, TiO2, FeO, SrO and BaO
1701
(mass%) analyses performed on prismatic Sr-barite (SB). A) and B) Macrocrystalline
1702
sphalerite (Sp) engulfing and replacing microcrystalline (MD) and blocky (BD)
1703
dolomite and partially filling vugular porosity (BSE); C) and D) Euhedral pyrite (Py),
1704
Sr-barite (SB), macrocrystalline calcite (MC) and prismatic macroquartz (MQ) filling
1705
vugular and fracture porosity in chert; E) and F) Vugular porosity completely cemented
1706
by prismatic Sr-barite (SB), macrocrystalline quartz (MQ), and macrocrystalline calcite
1707
(MC).
1708
Lithologic Types
Depth (m)
Comp.
Timing
FC
API hc (o)
Th (oC) range
Th (oC) average (n)
Recrystallized Calcite B Bioclast (CB)
br
XD26.0
hy
pr
wt
45-50
Recrystallized Fascicular Calcite (FC)
fcc fcc
XB93.5 XB93.5
aq hy
pr pr
wt
45-50
wt/bl wt
40-50 45-50
pr
wt
45-50
sec
wt
45-50
80-82 87-89 93-94 95-99 101-104 108.6-108.6 92-93 83-84 91-94 95-100 100-105 105.8-105.8 110-111 63-69 76-80 82-85 85-89 95-95 112.2-112.2 63-70 71-74 77.5-77.5 81-84 86-89 91.9-91.9 97.4-97.4 101-101 105-108 106-108 109-114 117-118 119-124 125-130 112-115 120-121
81.1 (2) 88.1 (5) 93.5 (2) 96.9 (5) 102.5 (5) 108.6 (1) 92.2 (2) 83.4 (2) 92.6 (4) 97.4 (8) 102.5 (3) 105.8 (1) 110.8 (2) 66 (2) 77.7 (3) 83.5 (6) 86 (5) 95 (1) 112.2 (1) 65.8 (8) 72.1 (3) 77.5 (1) 82.4 (4) 87.1 (2) 91.9 (1) 97.4 (1) 101 (1) 106.5 (2) 107 (3) 111.5 (5) 117.5 (2) 121.5 (3) 127.5 (2) 113.5 (2) 120.5 (2)
Mineral Host
Well
B B
Recrystallized Calcite A Spherulite (CS)
Macrocrystalline Calcite (MC)
C
sc
XC46.0
sig
XB71.5
sig/fcc
XB71.5/ XB98.0
fcc fcc
XB98.5 XB98.0
fcc
XB98.5
hy
aq
pr
pr/psec
Tm ICE (oC) range
Sal (wt% NaCl eq.)
-26.8
23.9-24.3 (*)
-20.9 -16.6 to -17.0 -20.1 to -20.4 -15.4 to -19.4 -18.1 to -20.8 -14.9 to -16.6 -14.8 to -17.2 -11.1 to -13.3 -16.7 to -18.4
23.0 19.9-20.2 20.5-22.4 19.0-22.0 21.0-22.9 18.6-19.9 18.5-20.4 15.1-17.2 20.0-21.3
fcc fcc
XB98.0 XB98.5
sig fcc
XB71.5 XB98.0
fcc sig
XB98.5 XB71.5
fcc
XB98.5
ct
XB91.0
fcc ct sig ct/fcc ct
XB98.5 XB91.0 XB71.5 XB91.0/ XB98.5 XB91.0
A
sc
XC46.0
C A
sig sc
C A C
sig sc sig
C
Macrocrystalline Quartz (MQ)
Prismatic Sr-barite (SB)
C
sec
hy
pr sec
aq
pr
yl yl
N/A N/A mx mx
pr/psec
sec
wt
45-50
XB71.5 XC46.0
pyl wt
33-35 45-50
XB71.5 XC46.0 XB71.5
yl wt
32-34 45-50
hy
aq
pr
pr
pr/psec
N/A 99-103 106-106 126-126 100-100 69-73 81-81 122-124 105-105 109-109 111-111 114-115 119-119 137-139 114-114 120-123 127-127 130-135 133-133 105-106 106-109 115-122 119-121 129-134 65-69.4 70-74.4 75-75 83-83 87.3-87.3 90-94 99-99 100.9-100.9 122-127 125-130 144-144 152-152 N/A 125-130 135-135
N/A (1) 101 (3) 106 (1) 126 (1) 100 (1) 71 (2) 81 (1) 123 (2) 105 (1) 109 (1) 111 (1) 114.5 (2) 119 (1) 138 (2) 114 (1) 121.5 (2) 127 (1) 132.5 (3) 133 (1) 105.5 (3) 107.5 (3) 118.5 (5) 120 (2) 131.5 (3) 68.2 (7) 73.1 (12) 75 (1) 83 (1) 87.3 (1) 92.1 (2) 99 (1) 100.9 (1) 124.5 (3) 127.5 (2) 144 (1) 152 (1) N/A (1) 127.5 (3) 135 (1)
-20.7 -15.2 to -16.1 -15.9 -16.0
22.9 18.8-19.5 19.4 19.5
-9.5 -9.0 -17.6 -15.4 to -15.9 -18.6 -17.9 to -18.0 -19.0 -19.9 to -21.2 -16.2 -16.2 to -17.8 -22.1 -16.6 to -18.2 -19.1 to -21.9 -19.5 to -21.3 -21.7 to -22.3 -19.3 to -20.5
13.4 12.9 20.7 19.0-19.4 21.4 20.9-21.0 21.7 22.3-23.1 19.6 19.6-20.8 23.7 19.9-21.1 21.8-23.6 22.0-23.2 23.5-23.9 21.9-22.7
-12.8 to -18.2 N/A -15.3 -16.6 -19.1 -18.1 to -19.6 -18.9
16.7-21.1 N/A 18.9 19.9 21.8 21.0-22.1 21.6
sec C
Saddle Dolomite (SD) C
C
sig
XB71.5
hy
pr/psec sec
br fcc br fcc/br
aq
pr
br
XF39.8 XB98.0 XF39.8 XB98.0/ XF39.8 XF39.8
br
XF39.8
hy
pr pr/psec
yl yl
yl yl
32-34 mx
N/A mx
N/A 122-124 83-83 100-100 105-107 119-119 N/A N/A 105-105 121-124 126-131 132-137 137-137 140-143 144-148 150-150 85-85 79-79 89-89
N/A (1) 123 (2) 83 (1) 100 (1) 106 (2) 119 (1) N/A (2) N/A (7) 105 (1) 122.5 (3) 128.5 (4) 134.5 (6) 137 (1) 141.5 (2) 146 (2) 150 (1) 85 (1) 79 (1) 89 (1)
-18.2 -16.9 to -19.2
21.1 20.2-21.8
-16.6 to -17.3 -19.3 to -25.3 -16.1 -18.4 to -21.0 -13.5 to -20.7 -15.0 to -17.8 -16.7 -14.5 to -16.8 -16.9 to -17.4 -20.4
19.9-20.5 21.9-26.1 19.5 21.3-23.0 17.3-22.9 18.6-20.8 20.0 18.2-20.1 20.2-20.5 22.7
Mineral Phase
Sampling
Well Depth (m) δ13C
Microcrystalline/Blocky Calcite (MBCB) (Bioclastic Grainstone and Rudstone)
bulk
B
Microcrystalline/Blocky Calcite (MBCL) (Laminite)
bulk
B
Fascicular Calcite (FC)
bulk
C B
C
micromill
C
XD06.0 XD08.0 XD10.0 XD73.0 XE38.3 XC95.4 XD00.0 XD01.0 XD43.0 XD44.0 XA15.0 XA37.0 XC85.0 XC88.2 XC89.0 XC93.0 XB57.0 XA06.0 XA95.0 XC02.3 XC04.0 XB98.8 XC01.0 XC08.5 XB83.0 XB87.0 XB88.5
XB98.8 XC08.5 Calcite Spherulite (CS)
bulk
A
XA66.5 XA84.0
1.10 0.99 1.13 0.83 0.45 1.14 0.70 1.42 1.37 1.43 1.41 0.93 1.30 1.28 1.52 1.42 1.07 1.98 1.43 1.60 1.62 1.42 1.17 1.26 0.05 1.15 0.88 0.71 0.34 0.19 0.16 0.59 0.72 0.63 0.98 0.52
δ13C error (1s) 0.04 0.06 0.05 0.06 0.05 0.03 0.05 0.04 0.04 0.05 0.05 0.05 0.07 0.05 0.05 0.02 0.05 0.02 0.09 0.05 0.04 0.05 0.05 0.05 0.06 0.04 0.03 0.04 0.05 0.03 0.06 0.04 0.02 0.04 0.05 0.05
δ18O -0.95 1.04 0.55 -1.06 -0.80 -0.61 -0.54 0.46 -1.18 -0.90 -0.59 -0.19 -0.99 -0.97 -0.23 0.63 0.54 -2.94 -1.93 -3.94 -3.34 -4.97 -4.36 -2.34 -5.70 -5.25 -5.10 -4.95 -5.37 -5.24 -5.12 -5.19 -3.85 -2.88 1.23 1.28
δ18O error (1s) 0.06 0.05 0.08 0.10 0.09 0.06 0.08 0.05 0.04 0.08 0.11 0.08 0.08 0.06 0.06 0.04 0.07 0.08 0.07 0.08 0.08 0.07 0.07 0.07 0.06 0.09 0.02 0.05 0.02 0.03 0.06 0.03 0.03 0.04 0.07 0.07
Sampling Observation Bulk analyses in the bioclastic grainstone composed of blocky and microcrystalline calcite in recrystallized bioclasts (>90%) and microcrystalline dolomite in partially dolomitized portions (<10%).
Bulk analyses in the bioclastic rudstone composed of blocky and microcrystalline calcite in recrystallized bioclasts (100%).
Bulk analyses in the laminite lithofacies constituted by recrystallized blocky and microcrystalline calcite (>90%) partially replaced by microcrystalline dolomite (<10%).
Bulk analyses in sample composed of syngenetic fascicular calcite crusts (>90%), eodiagenetic calcite spherulite (<10%) and eo/mesodiagenetic blocky dolomite (<10%).
Punctual analyses performed in the syngenetic fascicular calcite and/or fascicular aggregates.
Bulk analyses in sample constituted by eodiagenetic calcite spherulite (>90%), syngenetic fascicular calcite (<10%) and eo/mesodiagenetic blocky dolomite (<10%).
B
C
bulk
B
micromill
C
Macrocrystalline Calcite bulk (MC) micromill
C
Microcrystalline/Blocky Dolomite (MBDD) (Dolostone)
C
XB67.0 XA03.0 XA41.2 XA58.5 XA69.8 XB59.0 XC01.0 XC91.0 XA85.0 XB53.0 XD51.0 XC87.4 XC94.0 XC08.5 XB83.0 XB84.0 XB71.5 XB83.0 XB87.0 XB88.5 XC08.5
Saddle Dolomite (SD)
bulk
A
B C
XC88.5 XC96.6 XD19.0 XD19.5 XE62.0 XF17.0 XD90.0
2.43 1.25 0.75 1.79 1.49 2.14 1.54 1.91 0.31 1.36 1.67 1.25 1.16 1.44 1.37 1.28 1.18 -0.27 0.75 0.90 0.97 0.35 0.83 0.01 0.70 0.88 -1.88 -0.57 -0.04 0.22 1.03 0.34 -0.92
0.05 0.04 0.06 0.07 0.04 0.08 0.05 0.02 0.05 0.05 0.05 0.03 0.05 0.03 0.01 0.05 0.05 0.07 0.03 0.03 0.05 0.07 0.04 0.04 0.02 0.02 0.05 0.05 0.05 0.05 0.07 0.03 0.05
1.47 -2.04 0.76 -0.10 0.07 -0.56 -1.05 1.01 -2.11 -1.54 1.74 0.35 -1.12 -1.51 -1.45 -5.90 -4.92 -9.74 -9.51 -7.85 -6.19 -8.11 -7.67 -8.22 -7.82 -7.21 -6.87 -7.39 -9.49 -7.87 -7.33 -6.97 -6.35
0.07 0.07 0.10 0.08 0.07 0.12 0.09 0.08 0.07 0.07 0.07 0.10 0.07 0.04 0.04 0.07 0.07 0.07 0.03 0.04 0.05 0.07 0.04 0.05 0.02 0.02 0.07 0.07 0.07 0.07 0.07 0.11 0.07
dolomite (<10%).
Bulk analyses in the dolostone with blocky and microcrystalline dolomite (100%). Punctual analyses performed in the eo/mesodiagenetic blocky dolomite filling secondary porosity. Bulk analyses in sample composed of hydrothermal macrocrystalline calcite (>90%) and syngenetic fascicular calcite (<10%). Punctual analyses performed in the hydrothermal macrocrystalline calcite filling vugular and fracture porosity.
Bulk analyses in the dolostone and/or chert constituted by hydrothermal saddle dolomite (>90%) and macrocrystalline calcite (<10%) filling secondary porosity.
Mineral Phase
Sampling
Well Depth (m)
87
Fascicular Calcite (FC)
micromill
C
XB83.0 XB88.5 XB98.8 XC08.5
Calcite Spherulite (CS)
bulk
A
micromill
C
XA66.5 XA84.0 XB67.0 XB89.5
XB91.0
Sr/86Sr
± 2SE
Sampling Observation
0.71317 0.71329 0.71332 0.71360
0.00002 0.00002 0.00002 0.00001
Punctual analyses performed in the syngenetic fascicular calcite and/or fascicular aggregates.
0.71317 0.71338 0.71394 0.71317 0.71367 0.71365 0.71299 0.71360 0.71315
0.00001 0.00001 0.00001 0.00001 0.00001 0.00001 0.00002 0.00001 0.00001
Bulk analyses in sample constituted by eodiagenetic calcite spherulite (>90%), syngenetic fascicular calcite (<10%) and eo/mesodiagenetic blocky dolomite (<10%). Punctual analyses performed in the eodiagenetic calcite spherulite and spherulitic aggregates .
Microcrystalline/Blocky Dolomite (MBDD)
micromill
C
XC08.5
0.71351 0.71348
0.00001 0.00002
Punctual analyses performed in the eo/mesodiagenetic blocky dolomite filling secondary porosity.
Macrocrystalline Calcite (MC)
micromill
C
XB71.5
D
XD22.3
0.71136 0.71138 0.71199 0.71225 0.71215 0.71210
0.00005 0.00001 0.00001 0.00001 0.00001 0.00001
Punctual analyses performed in the hydrothermal macrocrystalline calcite (>90%) and Sr-barite (<10%) filling vugular and fracture porosity.
XD35.7 XD38.7 Saddle Dolomite (SD)
bulk
A
XC88.5 XC96.6 XD19.0 XD19.5
0.71109 0.71100 0.71123 0.71124
0.00001 0.00001 0.00001 0.00003
Bulk analyses in the dolostone and/or chert constituted by hydrothermal saddle dolomite (>90%) and macrocrystalline calcite (<10%) filling secondary porosity.
Sr-barite (SB)
micromill
D
XD35.7 XD41.6 XD42.1
0.71226 0.71209 0.71203 0.71202 0.71204
0.00001 0.00001 0.00001 0.00001 0.00001
Punctual analyses performed in the hydrothermal Sr-barite (>90%) and macrocrystalline calcite (<10%) filling vugular and fracture porosity.
Mineral Phase
δ18Ow Min
Max
δ18Ocarb Min
Fascicular Calcite (FC)
6.5
10.4
Fascicular Calcite (FC)
5.4
Calcite Spherulite (CS)
Max
Fluid Inclusion
Th (º C) Min
Max
-5.7
-1.9
Aqueous
92.2
92.2
12.4
-5.7
-1.9
Oil
83.0
111.0
6.3
16.3
-2.1
1.7
Oil
63.0
112.2
Saddle Dolomite (SD)
2.3
5.6
-9.5
-6.4
Aqueous
105.0
150.0
Macrocrystalline Calcite (MC)
3.3
11.1
-9.7
-4.6
Aqueous
101.0
130.0
Highlights A deep-burial hydrothermal system is recognized in the Campos Basin Pre-Salt carbonates. Hydrothermal phases show low δ18O values, and high Th and salinity fluid inclusions. The geochemical and fluid inclusion signatures are similar to MVT deposits. Hydrothermal fluids are the result of mixing between different sources. Atypical oil generation and migration occurred associated with the hydrothermal action.
Declaration of interests ☒ The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper. ☐The authors declare the following financial interests/personal relationships which may be considered as potential competing interests:
S0264-8172(19)30595-1
DOI:
https://doi.org/10.1016/j.marpetgeo.2019.104143
Reference:
JMPG 104143
To appear in:
Marine and Petroleum Geology
Received Date: 3 July 2019 Revised Date:
18 November 2019
Accepted Date: 18 November 2019
Please cite this article as: Lima, Bruno.Eustá.Moreira., Tedeschi, L.R., Pestilho, André.Luiz.Silva., Santos, R.V., Vazquez, J.C., Guzzo, J.V.P., De Ros, L.F., Deep-burial hydrothermal alteration of the Pre-Salt carbonate reservoirs from northern Campos Basin, offshore Brazil: Evidence from petrography, fluid inclusions, Sr, C and O isotopes, Marine and Petroleum Geology (2019), doi: https:// doi.org/10.1016/j.marpetgeo.2019.104143. This is a PDF file of an article that has undergone enhancements after acceptance, such as the addition of a cover page and metadata, and formatting for readability, but it is not yet the definitive version of record. This version will undergo additional copyediting, typesetting and review before it is published in its final form, but we are providing this version to give early visibility of the article. Please note that, during the production process, errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain. © 2019 Published by Elsevier Ltd.
1
1
Deep-burial hydrothermal alteration of the Pre-Salt carbonate
2
reservoirs from northern Campos Basin, offshore Brazil: Evidence
3
from petrography, fluid inclusions, Sr, C and O isotopes
4 a, b
, Leonardo Ribeiro Tedeschi c, André Luiz Silva
5
Bruno Eustáquio Moreira Lima
6
Pestilho c, Roberto Ventura Santos d, Joselito Cabral Vazquez c, Jarbas Vicente Poley
7
Guzzo c, Luiz Fernando De Ros b
8 9
a
Petrobras S.A., Avenida República do Chile, 330, East Tower, Rio de Janeiro, RJ,
10
20031-170, Brazil; [email protected]
11
b
12
Gonçalves, 9500, Porto Alegre, RS, 91509-900, Brazil; [email protected]
13
c
14
Universitária, Rio de Janeiro, RJ, 21941-915, Brazil;
15
[email protected]; [email protected];
16
[email protected]; [email protected]
17
d
18
Central, Brasília, DF, 70910-900, Brazil; [email protected]
19 20 21
Graduate Program in Geosciences, Rio Grande do Sul Federal University, Av. Bento
Petrobras, Research Center (CENPES), Avenida Horácio Macedo, 950, Cidade
Geosciences Institute, University of Brasília, Campus Universitário Darcy Ribeiro, Ala
2 22
ABSTRACT
23
Petrographic, mineralogical, elemental, isotopic and fluid inclusion analyses were
24
integrated to unravel the diagenetic evolution of Brazilian Pre-Salt lacustrine carbonate
25
reservoirs of northern Campos Basin, southeast Brazilian margin. Detailed thin section
26
and cathodoluminescence petrography, scanning electron microscopy and electron
27
microprobe analyses established a paragenetic evolution of diagenetic processes and
28
products, comprising extensive dolomitization, silicification, and dissolution. A
29
paragenesis including saddle dolomite, macrocrystalline calcite, mega-quartz, Sr-barite,
30
celestine, fluorite, dickite, sphalerite, galena, and other metallic sulfides filling fractures
31
and dissolution porosity, and aqueous fluid inclusions with homogenization
32
temperatures of 92 to 152 °C and salinities between 13 to 26 wt. % eq. NaCl
33
characterized a hydrothermal system with some analogy to carbonate-hosted Pb-Zn
34
Mississippi Valley (MVT) and Irish-type deposits. Petroleum inclusions and solid
35
bitumen testify atypical oil generation and migration, associated with the hydrothermal
36
flow. The host Pre-Salt spherulitic and fascicular carbonates present highly radiogenic
37
87
38
Hydrothermal phases show δ18O values more negative than syngenetic and diagenetic
39
carbonates. The δ13C values are interpreted as result of interaction between the
40
hydrothermal fluids and the host rocks. The combined data set provides clear evidence
41
of intense hydrothermal alteration of northern Campos Basin Pre-Salt reservoirs at
42
deep-burial conditions (> 2 km), possibly related to Late Cretaceous or more probably
43
Paleogene magmatic activity. Mixed-sourced fluids bearing a basinal signature fed the
44
hydrothermal system and promoted dissolution of the host rocks. The hydrothermal
45
alterations had strong impact on the porosity, permeability, and heterogeneity,
Sr/86Sr ratios, indicating strong interaction with continental crust materials.
3 46
contributing, together with the associated fracturing, to the excellent production
47
performance of the Pre-Salt reservoirs.
48
Keywords Pre-Salt; lacustrine carbonates; Campos Basin; deep-burial; hydrothermal
49
alteration; magmatic activity
50
1. Introduction
51
Carbonate reservoirs usually have complex and heterogeneous diagenetic
52
evolution as result of their reactive character. Common association with faults and
53
fracture systems often results in intense dolomitization, silicification and/or dissolution
54
(e.g., Gregg and Sibley, 1987; Machel and Mountjoy, 1987; Packard et al., 2001;
55
Corbella et al., 2006; Wilson et al., 2007; Callot et al., 2010; Fontana et al., 2010).
56
Hydrothermal alteration of carbonate rocks has been described in diverse basins, and its
57
characterization is important, both to differentiate from the “conventional” diagenetic
58
evolution of carbonates, and because hydrothermal processes substantially impact,
59
positively or negatively, the quality and performance of carbonate reservoirs (e.g.,
60
Machel and Lonnee, 2002; Davies, 2002, 2004; Machel, 2004; Davies and Smith, 2006;
61
Lonnee and Machel, 2006; Biehl et al., 2016).
62
The hydrothermal
alteration of carbonate deposits involves complex
63
physicochemical processes of interaction with hot fluids, promoting the precipitation of
64
mineral assemblages commonly including saddle dolomite, fluorite, barite, anhydrite,
65
dickite, sphalerite, and pyrite (Neilson and Oxtoby, 2008). These processes indicate the
66
existence of geothermal anomalies, as well as permeable pathways, such as deep fault
67
systems, to allow the percolation and interaction of fluids (Davies and Smith, 2006; Xu
68
et al., 2015; Mansurbeg et al., 2016).
4 69
The Lower Cretaceous Pre-Salt carbonate deposits of the Coqueiros and Macabu
70
formations of northern Campos Basin host important hydrocarbon reservoirs (Dias et al.,
71
1988; Winter et al., 2007). In Pre-Salt deposits there are significant variations in their
72
quality (porosity and permeability) due to the complex diagenesis (Herlinger Jr. et al.,
73
2017; Lima and De Ros, 2019). Recently, evidence on hydrothermal alteration of Pre-
74
Salt deposits has been presented for the Campos Basin (Alvarenga et al., 2016; Vieira
75
de Luca et al., 2017; Lepley et al., 2017; Herlinger Jr. et al., 2017; Lima and De Ros,
76
2019) and for Kwanza Basin, offshore Angola, its western African counterpart (Poros et
77
al., 2017; Girard and San Miguel, 2017; Teboul et al., 2017, 2019). Alvarenga et al.
78
(2016) identified hydrothermal vents in seismic sections of central Campos Basin.
79
Vieira de Luca et al. (2017), Poros et al. (2017), Girard and San Miguel (2017) and
80
Teboul et al. (2019) suggested that hydrothermal fluids ascribed pervasive silicification
81
of Pre-Salt carbonates based on petrographic, isotopic data and fluid inclusions.
82
Herlinger Jr. et al. (2017) and Lima and De Ros (2019) have shown evidence of
83
hydrothermal alteration in Pre-Salt reservoirs from northern Campos Basin, based on
84
the description of hydrothermal mineral phases in thin sections, such as coarse quartz,
85
saddle dolomite, sulfides and sulfates.
86
Although evidence of hydrothermal action in Campos Basin have been discussed
87
by Vieira de Luca et al. (2017), Herlinger Jr. et al. (2017) and Lima and De Ros (2019),
88
a detailed study of hydrothermal mineral phases is needed to improve the
89
characterization of the hydrothermal processes in this and other Pre-Salt areas, and has
90
not been published yet. Particularly, it is important to examine the fracture-related
91
dolomitization, silicification, dissolution, and the precipitation of typical hydrothermal
92
mineral phases, such as saddle dolomite, macrocrystalline calcite, mega-quartz, sulfides
93
and sulfates filling secondary porosity. With this aim, an integrated study was
5 94
performed on four cored wells, comprising detailed petrographic characterization
95
performed through microprobe, electron and optical microscopy, temperature and
96
salinity data from fluid inclusion analyses in diagenetic carbonates, quartz and Sr-barite,
97
and C, O and Sr isotopic analyses in carbonates and Sr-barite. Understanding the
98
processes and patterns of hydrothermal alteration of Pre-Salt deposits is of key
99
importance for reservoir quality assessment during exploration and for the production
100
101
optimization of these extraordinary reservoirs.
2. Geological setting
102
The Campos Basin is situated in southeastern Brazilian margin and covers an
103
area of approximately 100,000 km2, with only 5,800 km2 onshore. The study area is
104
located in the offshore part of the northern Campos Basin (Fig. 1). The stratigraphy of
105
the study area is based on the stratigraphic chart from Winter et al. (2007) (Fig. 2). The
106
reservoirs being addressed here are part of the Coqueiros (rift stage) and Macabu (sag
107
phase) formations from of the Pre-Salt interval of the basin.
108
The rift sequence lies on a basement of the Neoproterozoic Ribeira Fold Belt,
109
which includes a variety of metamorphic and magmatic rocks, such as tonalities,
110
granodiorites, granites and gabbros (Winter et al., 2007; Tupinambá et al., 2012). The
111
initial rift evolution of the basin is represented by Hauterivian volcanics from the
112
Cabiúnas Formation and by Barremian to Lower Aptian sediments of the Itabapoana,
113
Atafona, and Coqueiros formations (basal interval of the Lagoa Feia Group; Fig. 2). The
114
Itabapoana and Atafona formations consist of conglomerates, sandstones, siltstones,
115
arenites of stevensitic ooids and shales (Armelenti et al., 2016). The Coqueiros
116
Formation is composed by bioclastic rudstones and grainstones intercalated with
6 117
dolostones and organic-rich shales (Baumgarten et al., 1988; Dias et al., 1988; Castro,
118
2006; Thompson et al., 2015).
119
The post-rift is defined by the upper interval of the Lagoa Feia Group, during
120
which evolved the so-called sag stage of the basin. In addition to Itabapoana Formation,
121
this sequence comprises the Gargaú, Macabu, and Retiro formations that were deposited
122
during the Middle/Upper Aptian (Fig. 2). The Macabu Formation comprises
123
argillaceous and carbonate laminites as well as spherulitic and fibrous carbonate crusts.
124
These later strata are interpreted as chemical precipitates deposited in alkaline lacustrine
125
environments under arid climate conditions (Wright, 2011, 2012; Tosca and Wright,
126
2014; Wright and Barnett, 2014, 2015; Wright and Tosca, 2016; Herlinger Jr. et al.,
127
2017; Lima and De Ros, 2019).
128
The Retiro Formation (Fig. 2), a thick accumulation of evaporites, was deposited
129
during the late Aptian (Late Alagoas local stage) marine incursion, under arid climate
130
conditions (Leyden et al., 1976; Winter et al., 2007), being composed essentially of
131
anhydrite, halite, and also by bittern salts, such as sylvite and carnallite (Rodriguez et al.,
132
2018). These evaporites provide the stratigraphic seal for the large Pre-Salt hydrocarbon
133
accumulations, as well as for the migration of hydrothermal fluids through deep faults
134
(Lima and De Ros, 2019). Above the evaporites, a carbonate/clastic sequence was
135
deposited from Albian to Recent, which are out of the scope of the present study.
136
Magmatic activity has been recorded in different stages of Campos Basin
137
evolution. Winter et al. (2007) described important extrusive (basalts and hyaloclastites)
138
and intrusive (diabase and gabbro) magmatic events with Ar/Ar ages from 81.5 to 83.2
139
Ma. In the Paleogene, volcanic activity is represented by basaltic lava flows, dated 65.5
140
Ma (close to the Cretaceous-Paleogene limit), 62 Ma (Danian, Paleocene), 53 Ma
7 141
(Ypresian, Eocene), and alkaline basalts, diabase and volcanic tuffs at approximately 43
142
Ma (Lutetian, Eocene).
143
3. Analytical methods
144
More than 300 core plugs and sidewall core samples from four wells were taken
145
from pervasively dolomitized and silicified intervals of the Campos Basin Pre-Salt
146
rocks (Fig. 1). Oil and salt were removed from all samples, using toluene and methanol
147
in a Soxhlet apparatus. Standard-thickness thin sections were prepared from blue epoxy
148
resin-impregnated samples for transmitted light microscopy. All thin sections were
149
examined in petrographic microscopes under uncrossed (//P) and crossed polarizers
150
(XP).
151
The paragenetic evolution of diagenetic mineral phases was examined in
152
approximately forty polished thin sections using a combination of cathodoluminescence,
153
scanning electron microscopy and microprobe analyses. Cathodoluminescence (CL)
154
was performed with a CITL Mk5-2 equipment on a Zeiss Axiocam MRC microscope.
155
Scanning electron microscopy images with secondary electrons (SEM) and
156
backscattered electrons (BSE) were obtained with a JSM-IT300 JEOL electron
157
microscope. Mineral chemical composition was verified with the support of
158
wavelength-dispersive X-ray spectroscopy (WDS) in a JEOL JXA-8230 electron
159
microprobe, and energy-dispersive spectrometry (EDS) in an OXFORD X-MaxN
160
equipment. Mineralogical mapping was also performed by QEMSCAN 650 (FEI)
161
equipment. These analyses were performed at the Petrobras Research Center (CENPES)
162
and at the Regional Center for Technological Development and Innovation (CRTI) of
163
Goiás Federal University (UFG).
8 164
Fluid inclusion (FI) petrography and microthermometry was performed on eight
165
double-polished thin sections (120 µm-thick) prepared from sidewall core samples
166
following standard procedures (Goldstein and Reynolds, 1994). These sections were
167
selected to represent the lithologies and mineral phases of the rift and sag sections (red
168
squares in Fig. 3). FIs were analyzed in macrocrystalline quartz (MQ), Sr-barite (SB),
169
saddle dolomite (SD), blocky calcite in recrystallized bioclasts (MBCB), and
170
macrocrystalline (MC), fascicular (FC) and spherulitic (CS) calcite as described in the
171
results section. A total of 203 determinations of homogenization temperature (Th) and
172
ice melting temperature (Tm ICE) of fluid inclusions hosted in syngenetic, diagenetic and
173
hydrothermal mineral phases were obtained using a Fluid Inc. USGS gas flow and a
174
Linkam THMSG600 heating and freezing stages (Table 1), under plane polarized light
175
and ultraviolet (UV; 380 nm) illumination.
176
FIs where described using the approach of Roedder (1984), where FIs where
177
classified according to its genetic type (primary, secondary, and pseudo-secondary) and
178
tied-in to the paragenetic sequence. FI types were characterized according to their
179
location, relationship to the host mineral, and consistency of visual parameters (e.g.,
180
apparent liquid/vapor ratio). Stage calibration was carried out using H2O-CO2 and pure
181
H2O-critical density synthetic inclusions. Homogenization and melting temperatures
182
were measured with a precision of ±0.1 °C. Minimum salinities were estimated using
183
the ice melting temperatures computed from NaCl-H2O system (Bodnar and Vityk,
184
1994). Salinities of inclusions displaying CaCl2-NaCl-H2O eutectic temperatures (-
185
52 °C) were estimated by combining of hydrohalite and ice melting temperatures
186
(Steele-MacInnis et al., 2011). For petroleum inclusions, API gravity was estimated
187
with a microthermometric technique, based on the oil fluorescence results (Goldstein
188
and Reynolds, 1994).
9 189
A total of 69 carbon (δ13C) and oxygen (δ18O) isotopes analyses were performed
190
on 51 samples derived from core plugs and sidewall cores of three wells (A, B and C)
191
(Table
192
microcrystalline/blocky calcite from bioclasts (MBCB) and laminites (MBCL),
193
microcrystalline/blocky dolomite from dolostones (MBDD), saddle dolomite (SD), and
194
macrocrystalline (MC), fascicular (FC) and spherulitic (CS) calcite. Forty-eight bulk
195
rock analyses were obtained from samples with more than 90% of a single carbonate
196
phase, using an IRMS Thermo Scientific Delta XL Advantage coupled to a Gasbench II
197
device in the University of São Paulo (USP) Lab. In addition, 21 samples (~100 mg) of
198
selected carbonate phases were obtained for carbon and oxygen isotopes from eight
199
polished thin sections (120 µm-thick), using a computer-monitored New Wave
200
Research MicroMill TM equipment, and a Thermo Scientific Delta V Isotope Ratio
201
Mass Spectrometer (IRMS), coupled to a Gasbench II unit in CENPES. All data were
202
reported relative to the Vienna Pee Dee Belemnite (V-PDB) international standard in
203
per mil units (‰). External precision and accuracy were checked in relation to the
204
reference material NBS 19 (TS-Limestone; Coplen et al., 2006), yielding results better
205
than ±0.09‰ for δ13C and ±0.10‰ for δ18O for both laboratories.
2;
blue
circles
The analyses of
206
87
in
Fig.
3).
The
analyzed
mineral
phases
were
Sr/86Sr ratios of thirty samples were obtained at the
207
Geochronology Lab of the University of Brasília (UnB). Roughly 1 mg of carbonate
208
powder was placed into SavillexTM (PFA tubes) and digested using an 0.5 mol/L acetic
209
acid solution in order to avoid leaching of non-carbonate phases. After centrifugation
210
the supernatant was dried in a hot plate and dissolved in nitric acid (2.9 mol/L). Sr was
211
separated from Rb in Teflon chromatographic columns filled with SrSpec (100-150 µm)
212
resin.
213
mass spectrometer in static multicollector (TIMS). The Sr ratios were normalized
87
Sr/86Sr ratios were determined using a Thermo Triton Plus thermal ionization
10 214
assuming a 88Sr/86Sr ratio of 0.1194. External reproducibility of 87Sr/86Sr ratios from the
215
NBS-987 standard yielded 0.710271 ± 0.000017 (n=26). Sr isotope analyses were
216
normalized to NBS-987 value of 0.710250.
217
4. Results
218
4.1. Petrography
219
Bioclastic grainstones and rudstones, stevensitic ooidal arenites (stevensite is a
220
smectite group clay mineral), and dolostones are the main lithologic types in the rift
221
interval (Herlinger Jr. et al., 2017; Lima and De Ros, 2019; Fig. 4A and B). Bivalves,
222
ostracods, and gastropods are the main bioclasts in the grainstones and rudstones.
223
Common diagenetic processes in these rocks include bioclast micritization,
224
neomorphism, recrystallization, replacement and dissolution. Most of the bivalves occur
225
recrystallized, and/or dissolved and filled by diagenetic cements. Cementation of inter-
226
and intraparticle, intercrystalline, moldic, vugular, channel and fracture porosity types
227
occurs mainly by drusy/rim pore-lining calcite, blocky and saddle dolomite, fibrous
228
chalcedony, microcrystalline, prismatic to coarse mosaic quartz, radiated and prismatic
229
celestine and Sr-barite, pyrite, svanbergite and bitumen. The bioclasts and the
230
stevensitic ooids are heterogeneously replaced by dolomite, calcite, microcrystalline
231
quartz, fibrous chalcedony, sulfides and svanbergite.
232
Lithologies described in the sag section include fascicular calcite crusts,
233
stevensitic claystones with calcite spherulites, intraclastic rudstones and grainstones,
234
laminites, dolostones, and cherts (Herlinger Jr. et al., 2017; Lima and De Ros, 2019; Fig.
235
4C to 4F). As for the rift bioclasts, most spherulitic and fascicular calcites are
236
recrystallized and/or replaced by silica or dolomite. The cherts consist of micro- and
237
macrocrystalline quartz, and spheroidal, fibrous chalcedony replacing syngenetic and
11 238
diagenetic constituents and filling pores. Rift lithologic types present a diagenetic
239
sequence similar to that of sag rocks. Therefore, we represent all diagenetic sequence by
240
using Figure 5, which shows the paragenetic sequence interpreted for the sag fascicular
241
calcite crusts, including 28 syngenetic, eodiagenetic, mesodiagenetic and hydrothermal
242
processes and products, schematically represented in Figure 6 (for more details in Lima
243
and De Ros, 2019).
244
Deposits of stevensite and other Mg silicates (event 1 in Fig. 5) occur throughout
245
the whole Aptian Pre-Salt succession, forming the substrate to the precipitation of
246
fascicular calcite crusts and spherulites (events 2 and 3 in Fig. 5) in a hyper-alkaline sag
247
lacustrine environment (Fig. 6A). Fascicular calcite crusts occur intercalated with
248
discontinuous levels of clay laminations, clay peloids and ooids of and siliciclastic
249
grains, such as muscovite and quartz, which were partially to totally replaced by
250
dolomite, calcite, microquartz, chalcedony and pyrite. Calcite spherulites were formed
251
as concretions, replacing and/or displacing these fine-grained deposits. Transition forms
252
between spherulitic and fascicular aggregates are common, including asymmetrical
253
spherulites, vertically elongated spherulites with lobate borders, hemi-spherulites, and
254
single or multiple fascicular aggregates nucleated on spherulites.
255
Sag lithologic types often occur partially to fully recrystallized, dolomitized,
256
silicified, cemented and/or dissolved. The Mg clay deposits occur frequently dissolved,
257
and/or replaced by silica, calcite, dolomite, pyrite or magnesite (events 4–7 in Fig. 5;
258
Fig. 6B). Lamellar magnesite replaced pseudomorphically the clay laminations in
259
subordinate amounts (event 6 in Fig. 5). Blocky dolomite, microcrystalline silica, quartz
260
and chalcedony commonly replaced the Mg clay matrix, calcite spherulites and
261
fascicular aggregates, and fill primary and secondary pores (events 8–10 in Fig. 5; Fig.
262
6B, C). Recrystallization, fracturing and pressure dissolution of fascicular and
12 263
spherulitic calcite was heterogeneous (events 11–15 in Fig. 5; Fig. 6C, D). Stylolites
264
occur
265
concentrations.
frequently associated
with
dolomitized
intervals
and
residual
pyrite
266
In some dolomitized or silicified samples, recognizing the original carbonate
267
fabrics modified by these processes is difficult, if not impossible. Dolomitization and
268
silicification are pervasive as disseminated, replacive, and filling cavities and fractures
269
associated with hydrothermal features and minerals (Figs. 7, 8). The common sag
270
replacive and pore-filling phases are cryptocrystalline to microcrystalline silica (events
271
10, 17 and 21 in Fig. 5; Fig. 7C, E); fibrous, spherulitic chalcedony (events 10 and 17 in
272
Fig. 5; Fig. 7D, E); mosaic, bladed and prismatic rims to drusiform calcite (events 12,
273
16 and 23 in Fig. 5; Figs. 7B, 8A, B, C, E, F); blocky, rims, mosaic, and saddle
274
dolomite (events 17 and 20 in Fig. 5; Figs. 7A, B, C, 8B, F); and macrocrystalline,
275
flamboyant and prismatic rims to drusiform quartz (events 21 and 22 in Fig. 5; Fig. 7D,
276
E, F).
277
Partial to total diagenetic recrystallization, dolomitization and/or silicification of
278
fascicular and spherulitic calcite are common (events 16–17 in Fig. 5; Fig. 6D, 8F). Late,
279
coarse, subhedral saddle or baroque dolomite, mega-quartz, and macrocrystalline calcite
280
(events 20, 22 and 23 in Fig. 5; Figs. 7, 8) is widely observed filling vugs and veins.
281
Saddle dolomite exhibits a characteristic xenotopic-C texture with curved crystal faces
282
and sweeping extinction (Gregg and Sibley, 1984; Davies and Smith, 2006; Fig. 7A).
283
The pervasive hydrothermal silicification (events 17, 21 and 22 in Fig 5; Fig. 6D, E, F)
284
is texturally distinct from the eodiagenetic silica crusts (event 10; Fig. 6C), which are
285
concordant
286
dolomitization. A complex evolution in pore fluid chemistry can be interpreted for the
287
hydrothermal alterations of Pre-Salt carbonate rocks.
with
depositional
bedding,
and
clearly occurred
subsequent
to
13 288
Svanbergite (SrAl3(PO4)(SO4)(OH)6) 1 to 100 µm crystals (event 18 in Fig. 5)
289
occur disseminated in microcrystalline silica and/or in chalcedony spherulites,
290
dominantly where silicification affected clay-rich protoliths. Radiated, prismatic and
291
oriented Sr-barite and celestine (event 24 in Fig. 5; Figs. 7B, F, 8A, C, D); fibro-
292
radiated zeolite (event 25 in Fig. 5; Fig. 8E); micro- to macrocrystalline, blocky and
293
prismatic sulfides (pyrite, sphalerite, chalcopyrite and galena; event 26 in Fig. 5; Fig.
294
7F); microcrystalline rutile associated with traces of native copper and zinc (event 27 in
295
Fig. 5); and massive bitumen with shrinkage cracks (event 28 in Fig. 5; Fig. 8F) are less
296
common. Sulfates, sulfides, zeolite and rutile occur filling vugular and fracture porosity,
297
often replacing dolomite, calcite and quartz. Sulfides are dominantly pyrite with lesser
298
sphalerite and minor chalcopyrite and galena. Fluorite and dickite were also described
299
in other wells of the area (Herlinger Jr. et al., 2017), and dawsonite was identified in the
300
XRD analyses (Lima and De Ros, 2019).
301
The pore types that occur in the sag lithologic types are growth-framework,
302
inter-crystalline aggregates, from Mg-clay dissolution, intercrystalline and intra-
303
crystalline aggregates, vugular, and fracture (Lima and De Ros, 2019). Important
304
fracturing and dissolution events occurred, which largely modified the pre-existing pore
305
systems (events 7, 13, 14, 19 and 21 in Fig. 5; Fig. 6). The hydrothermal alterations
306
were related to intense fracturing and dissolution processes (events 14, 19 and 21; Fig.
307
6D, E and F). The late hydrothermal saddle dolomite, macrocrystalline quartz and
308
calcite, sulfates, sulfides, zeolite, svanbergite, rutile, and bitumen filled growth
309
framework, and mostly fracture and dissolution pores (events 18–28 in Fig 5; Fig. 6D, E,
310
F). A later overprint of calcite-quartz-sulfides veins and breccias is also observed.
311
In general, microcrystalline (MD), blocky (BD) and saddle dolomite (SD)
312
present higher concentrations of Na2O, Al2O3, SiO2, MnO and FeO compared to
14 313
fascicular (FC), spherulitic (CS) and macrocrystalline calcite (MC) (WDS analysis; see
314
Supplementary material). SD shows higher values of SiO2 and lower MnO and FeO
315
than other dolomite types. Similarly, hydrothermal MC also has lower concentrations of
316
SiO2, SrO and BaO than FC and CS. In addition, hydrothermal sphalerite (Sp) shows
317
high Fe and Pb and low As, Co, Cu and Ni contents in WDS analyses (Supplementary
318
material). Hydrothermal euhedral pyrite (Py) filling secondary porosity also occurs with
319
high Pb and low As, Co, Cu, Zn and Ni concentrations (WDS). Hydrothermal prismatic
320
Sr-barite (SB) presents SrO content ranging from 2% to 16%, with an average of 6.7%,
321
high TiO2 and low CaO and FeO concentrations (WDS).
322
4.2. Fluid inclusion analyses
323
Aqueous and petroleum fluid inclusions (FIs) hosted in syngenetic and
324
diagenetic/hydrothermal constituents (Fig. 9) were analyzed in the main rift and sag
325
lithologic types (Figs. 4, 7, 8), such as fascicular calcite crusts, stevensitic claystones
326
with calcite spherulites, intraclastic grainstones, bioclastic rudstones and cherts. The
327
phases analyzed were recrystallized bioclasts, fascicular and spherulitic calcite, and
328
diagenetic/hydrothermal macrocrystalline calcite, saddle dolomite, mega-quartz, and
329
prismatic Sr-barite, replacing pre-existing constituents, filling fractures and dissolution
330
porosity. Fluid inclusions (both aqueous and petroleum inclusions) are trapped in
331
recrystallized fascicular and spherulitic calcite displaying randomly distribution in areas
332
with clear mineral substitution. On the other hand, the preserved carbonate elements do
333
not host petroleum inclusions. Therefore, these inclusions in the recrystallized
334
carbonates
335
microthermometric results are displayed in Table 1.
were
interpreted
as
primary.
Petrographic
observations
and
15 336
All aqueous inclusions are two-phase (liquid plus vapor) at room temperature
337
without large variation in the vapor fraction, with exception of some primary inclusions
338
in saddle dolomite and macrocrystalline calcite displaying necking-down features,
339
which were disregarded from subsequent interpretations. Six high Th results from
340
saddle dolomite (140, 144, 148 and 150 °C) and Sr-barite (144 and 152 °C) may
341
represent stretched inclusions or unrecognized necking after a phase change (Goldstein,
342
2001).
343
Analyzed aqueous FIs were hosted in the following mineral phases:
344
recrystallized fascicular calcite (n=2), mega-quartz (n=34), prismatic Sr-barite (n=20),
345
macrocrystalline calcite (n=36), and saddle dolomite (n=34). The majority of the FIs
346
observed were primary and pseudo-secondary aqueous and petroleum types (78%), with
347
subordinate secondary inclusions (22%; Table 1). Aqueous secondary FIs are more
348
significant in macrocrystalline quartz (38%). The salinities of aqueous inclusions varied
349
between 12 and 26 wt. % eq. NaCl (Table 1).
350
Petroleum inclusion from sag samples correspond predominantly to white-blue-
351
fluorescent inclusions (n=82) and less yellow-fluorescent inclusions (n=16). Most are
352
hosted in the recrystallized fascicular and spherulitic calcite, and in the macrocrystalline
353
quartz, and less in macrocrystalline calcite, Sr-barite, and saddle dolomite (Table 1).
354
Petroleum FIs in rift samples consist of mostly white-blue-fluorescent (n=20) in
355
recrystallized calcite bioclasts, and less of yellow-fluorescent (n=3) in saddle dolomite,
356
macrocrystalline quartz and Sr-barite. Estimated API gravities range from 32 to 35° for
357
the yellow-fluorescent petroleum inclusions, and between 40 and 50° for the white-
358
blue-fluorescent inclusions. The characteristics of the FIs in specific analyzed
359
constituents can be summarized as follows:
16 360
Recrystallized calcite bioclasts (CB): 20 analyses of FIs in blocky calcite in
361
recrystallized calcite bioclasts (CB) from a rift rudstone in well B (Table 1; Figs. 3,
362
10B). Only primary petroleum FIs with white fluorescence (~45 to 50° API) were
363
observed. The Th values vary from 80 to 108.6 °C, with mean and median of 94.8 and
364
96.9 °C.
365
Recrystallized fascicular calcite (FC): 22 FI analyses in a sag recrystallized
366
syngenetic fascicular calcite (FC) sample from well B (Table 1; Figs. 3, 10C). Two
367
aqueous FIs present Th of 92 and 93 °C, and salinity of 23.9 and 24.3 wt. % eq. NaCl
368
(Fig. 11). White to blue fluorescent petroleum FIs (~40 to 50° API) have Th ranging
369
from 83 to 111 °C.
370
Recrystallized spherulitic calcite (CS): 38 FIs were analyzed in recrystallized
371
eodiagenetic calcite spherulite (CS) in a sag stevensitic claystone from well A (Table 1;
372
Figs. 3, 10D, E). Both primary and secondary petroleum inclusions with white
373
fluorescence color (~45 to 50° API) were hosted in the CS. Th of 18 primary petroleum
374
inclusions varies from 63 to 112.2 °C, with mean and mode of 83.5 °C (Fig. 10D). Th of
375
20 secondary petroleum inclusions ranges from 63 to 97.4 °C, with mean of 75.7 and
376
median of 72.1 °C.
377
Macrocrystalline calcite (MC): 36 primary and secondary aqueous and
378
petroleum FIs were measured in macrocrystalline calcite (MC) filling dissolution
379
porosity of three sag samples from well C (Table 1; Figs. 3, 12G, H). The 22 primary
380
and pseudo-secondary aqueous FIs present Th varying from 101 to 130 °C, with mean
381
of 114.3 and mode of 111.5 °C. Salinity of primary FIs ranges from 15.1 to 22.9 wt. %
382
eq. NaCl (Fig. 11A). Th of 5 secondary aqueous FIs varies from 99 to 126 °C, with
383
mean of 107 and mode of 101 °C (Fig. 12H). Salinity of secondary aqueous FIs ranges
384
from 18.8 to 22.9 wt. % eq. NaCl (Fig. 11B). Yellow-fluorescent (~30 to 35 °API)
17 385
primary inclusions have Th value of 100 °C. Th in 5 secondary petroleum FIs range
386
from 69 to 124 °C, with mean of 93.8 and median of 81 °C.
387
Macrocrystalline quartz (MQ): 58 FIs were observed in pore-filling mega-quartz
388
(MQ) in a silicified intraclastic grainstone, a FC crust, a chert and a stevensitic
389
claystone with CS from the sag phase in wells A and C (Table 1; Figs. 3, 12E, F). 16
390
primary/pseudo-secondary aqueous FIs display Th from 105 to 139 °C, with mean of
391
122.7 and mode of 132.5 °C (Fig. 12E). Salinity of these inclusions ranges from 12.9 to
392
23.7 wt. % eq. NaCl (Fig. 11A). Th of secondary aqueous FIs varies from 105 to 134 °C,
393
with mean of 116.6 and mode of 118.5 °C (Fig. 12F). Salinity of secondary aqueous
394
inclusions ranges from 19.9 to 23.9 wt. % eq. NaCl (Fig. 11B). Th of 26 white to yellow
395
fluorescent petroleum inclusions (~32 to 50° API) varies from 65 to 100.9 °C, with
396
mean of 76.3 and mode of 73.1 °C (Fig. 12E).
397
Prismatic Sr-barite (SB): 20 FIs analyses in prismatic Sr-barite (SB) cement
398
filling secondary porosity in a sag silicified intraclastic grainstone in well C (Table 1;
399
Figs. 3, 12C, D). Th of 11 primary/pseudo-secondary aqueous FIs ranges from 122 to
400
152 °C, with mean of 131.1 and mode of 127.5 °C (Fig. 12C). Th of 3 secondary
401
aqueous FIs varies from 122 to 124 °C (Fig. 12D). Salinity ranges from 16.7 to 22.1
402
wt. % eq. NaCl in 10 primary aqueous (Fig. 11A), and 20.2 to 21.8 wt. % eq. NaCl in 2
403
secondary aqueous FIs (Fig. 11B). Both primary and secondary petroleum FIs show
404
yellow fluorescence colors (~32 to 34° API). Only one primary yellow-fluorescent
405
petroleum inclusion yielded a Th value of 83 °C (Fig. 12C). Secondary petroleum FIs
406
have Th ranging from 100 to 119 °C with mean of 107.8 and mode of 106 °C (Fig. 12B).
407
Saddle dolomite (SD): 34 FIs in pore-filling saddle dolomite (SD) in sag FC
408
crusts and rift bioclastic rudstones from well C (Table 1; Figs. 3, 12I). Th measured in
409
20 primary aqueous FIs in zoned saddle dolomite varies from 105 to 150 °C with mean
18 410
of 132.8 and mode of 134.5 °C (Fig. 12I). Salinity of 29 primary aqueous FIs in SD
411
ranges from 16.9 to 26.1 wt. % eq. NaCl (Fig. 11A). Th obtained in 3 yellow-
412
fluorescent primary petroleum inclusions varies from 79 to 89 °C (Fig. 12I).
413
The salinities of inclusions hosted in recrystallized FC are close to the upper
414
limit for pore-filling hydrothermal minerals (Fig. 11A), with 97% higher than 15 wt. %
415
eq. NaCl. In the rift interval, salinities range from 17.3 to 22.7 wt. % eq. NaCl. The final
416
melting temperature of aqueous FIs hosted in SD, MC, MQ and SB ranges of -9 to -
417
25.3 °C, indicating NaCl-dominant compositions comparable to CaCl2 (e.g., Roedder
418
and Bodnar, 1997; Wilkinson, 2010).
419
Th of white-blue-fluorescent primary petroleum inclusions hosted in the
420
recrystallized CB, CS and FC ranges from 63 to 112.2 °C (Table 1; Fig. 10). Th of
421
yellow-fluorescent inclusions hosted in MC, MQ, SB and SD varies between 69 and
422
124 °C. Some yellow-fluorescent oil inclusions are interpreted to have been emplaced at
423
105 to 108 °C in MC and at 122 to 130 °C in SB in the presence of hypersaline aqueous
424
fluids. Mixed, co-genetic oil-aqueous FIs with Th ranging from 79 to 124 °C were
425
observed in fascicular calcite crusts, silicified intraclastic grainstone and bioclastic
426
rudstone (Table 1), indicating a common origin from a single fluid. Aqueous inclusions
427
coexist with yellow-fluorescent petroleum inclusions in MC, SB and SD.
428
4.3. Carbon and oxygen isotopes
429
The range of δ13C values for all analyzed carbonates is quite narrow, varying
430
between -1.88 and 2.43‰ (Table 2; Fig. 13A, B). The range of δ13C values for the
431
seven different carbonate phases are: 1) 0.45 to 1.43‰ for microcrystalline/blocky
432
calcite
433
microcrystalline/blocky calcite from laminites (MBCL); 3) 1.16 to 1.44‰ for
from
bioclasts
(MBCB);
2)
0.93
to
1.52‰
for
eodiagenetic
19 434
mesodiagenetic microcrystalline/blocky dolomite from dolostones (MBDD); 4) 0.31 to
435
2.43‰ for recrystallized eodiagenetic calcite spherulites (CS); 5) 0.05 to 1.98‰ for
436
recrystallized syngenetic fascicular calcite (FC); 6) -0.27 to 1.28‰ for diagenetic
437
macrocrystalline calcite (MC); and, 7) -1.88 to 1.03‰ for saddle dolomite (SD).
438
In contrast to the δ13C values, the δ18O values from the analyzed carbonates
439
show a wider variation range (from -9.74 to 1.74‰) (Table 2; Fig. 13A, B). The range
440
of δ18O values from the seven different carbonate phases are: 1) -1.18 to 1.04‰ for
441
MBCB; 2) -0.99 to 0.63‰ for MBCL; 3) -1.51 to 0.35‰ for MBDD; 4) -2.11 to 1.74‰
442
for CS; 5) -5.70 to -1.93‰ for FC; 6) -9.74 to -4.92‰ for MC; and, 7) -9.49 to -6.35‰
443
for SD. Samples analyzed in this study may be arranged into two main groups based on
444
the δ18O values: samples (MC and SD) with more depleted oxygen isotope values (<-
445
4.92‰); and samples (MBCB, MBCL, MBDD and CS) with more enriched oxygen
446
isotope values (>-2.11‰) (Fig. 13A). The δ18O values of FC samples (-5.70 to -1.93‰)
447
fall in between these two groups (Fig. 13A). The comparison of Fig. 13A and 13B with
448
13C and 13D show that both δ13C and δ18O values from different mineral phases are not
449
biased by different wells. Likewise, the distribution of samples against stratigraphy (Fig.
450
3) shows that sampling for isotope analyses are not biased against stratigraphy.
451 452
4.4. Strontium isotopes The analyzed samples show a wide range of
87
Sr/86Sr values (0.71100 to
453
0.71394; Table 3; Fig. 13). They may be separated into two main groups: those with Sr
454
isotope ratios above 0.71299 (MBDD, CS, FC, MC) and those below 0.71226 (MC, SD,
455
SB). As shown in Table 3, the Sr isotope values have a narrow range within each group
456
of samples. For instance, SD values range between 0.71100 and 0.71124, while CS
457
samples range between 0.71299 and 0.71394. As in the δ13C and δ18O results, 87Sr/86Sr
20 458
ratios from different mineral phases are not biased against stratigraphy (Fig. 3) or by
459
different wells (Fig. 13).
460
5. Discussion
461
5.1. Fluid temperature and salinity
462
There is no correlation between the salinity and temperature of fluid inclusions
463
(Fig. 14). Except for few FIs in macrocrystalline quartz and calcite, all other phases
464
exhibit similar salinity range. In contrast, the average Th of saddle dolomite and Sr-
465
barite inclusions are higher than those measured in the mega-quartz and calcite. Based
466
on these observations, the following points may be considered: i) the salinity variations
467
do not indicate mixing of fluids, but are related to cooling and/or neutralization by fluid-
468
rock interaction; ii) the formation of mega-quartz, saddle dolomite, macrocrystalline
469
calcite and Sr-barite, which are associated with higher temperature fluids, postdated the
470
formation of the other phases. This further indicates that heat was not pervasive in the
471
system, but rather concentrated along structures such as fractures that acted as conduits
472
for the hot fluids.
473
The variability of the Th values in aqueous inclusions may be related to post-
474
entrapment modification of FIs at variable high temperatures (King and Goldstein,
475
2018). For instance, the necking-down process occurs most rapidly at higher
476
temperatures, often producing Th values without any relation to the original
477
temperatures (Roedder, 1984; Bodnar et al., 1985; Goldstein, 2001; Bodnar, 2003). The
478
similar Th values in macrocrystalline calcite and quartz indicate that there were no
479
significant post-entrapment changes in FIs. Another relevant indication of preservation
480
of the integrity of the FIs is the robust petrographic control evidenced by the relatively
481
narrow range of Th values of individual fluid inclusions, generally much narrower than
21 482
within the entire sample, or between different samples (e.g., Wilkinson, 2010).
483
Therefore, the wider dispersion in Th values observed in hydrothermal saddle dolomite
484
is probably representing different generations of FIs trapped in the zoned crystals (Fig.
485
9D).
486
The Th values of aqueous FIs hosted in minerals precipitated or recrystallized by
487
hydrothermal fluids show a relatively narrow range, varying from to 92 and 152 °C
488
(Table 1; Fig. 14), in which 95% of the results range from 100 to 140 °C. In general,
489
there is a remarkable Th overlapping among all analyzed phases, with slightly higher
490
mean values for Sr-barite and saddle dolomite. The FI data indicate that hydrothermal
491
saddle dolomite, Sr-barite, macrocrystalline calcite, and mega-quartz precipitated over a
492
similar temperature range. The only two analyses performed in aqueous FIs of
493
recrystallized fascicular calcite yielded Th near the lower limits of the hydrothermal
494
phases (Table 1; Fig. 14). Poros et al. (2017) verified a Th variation between 95 to
495
115 °C in inclusions hosted in dolomite cement, 95 to 125 °C in silica (chert and
496
chalcedony) replacing carbonate, and 105 to 125 °C in pore-filling mega-quartz in the
497
Pre-Salt of Kwanza Basin, Angola. Although those mineral phases were interpreted to
498
have been formed by a hydrothermal system in a deep-burial condition, the Th values in
499
Kwanza Basin Pre-Salt reservoirs are somewhat lower than those obtained in Campos
500
Basin.
501
All analyzed phases in the study area contain abundant oil and/or co-genetic oil-
502
aqueous inclusions, which is indicative of atypical petroleum generation and migration
503
(sensu Magoon and Dow, 1994) associated with percolation of hydrothermal fluids
504
responsible for the recrystallization and precipitation of the host mineral phases. When a
505
sedimentary sequence is percolated by hot fluids, the organic matter of host rocks can
506
be thermally altered, generating petroleum through hydrothermal or “forced” maturation
22 507
(Ilchik et al., 1986; Anderson, 1991; Davies and Smith, 2006). The liquid petroleum
508
droplets produced by this mechanism are transported by the convecting fluids
509
(Schoenherr et al., 2007), and can be either trapped as petroleum FIs in hydrothermal
510
minerals, or devolatilized in the pore space as bitumen. This type of natural hydrous
511
pyrolysis is not efficient enough to fill a reservoir or an oil field (Davies and Smith,
512
2006) but may, however, generate oil or co-genetic oil-aqueous FIs and solid bitumen
513
(Simoneit, 1990; Simoneit, 2018, and references therein) comparable to those described
514
here, which may be associated to hydrothermal dolomitization and ore deposits
515
described in diverse geotectonic settings (Disnar, 1996; Wilson and Zentilli, 1999;
516
Parnell et al., 2003; Bertrand et al., 2003; Neilson and Oxtoby, 2008; Suchý et al., 2010;
517
Greenwood et al., 2013 and references therein; Ostendorf et al., 2015).
518
With progressive burial, ‘normal’ background mesodiagenetic processes
519
proceeded in Campos Basin Pre-Salt independently of hydrothermal events (Fig. 5).
520
Thus, it should be noted that part of the petroleum inclusions hosted in the recrystallized
521
calcite bioclasts and spherulites (Th = 80–111 °C), in which no aqueous inclusions were
522
observed, may have been trapped during the migration of oil generated by the
523
conventional petroleum system. According to Mello et al. (1994), conventional burial
524
petroleum generation, migration and accumulation started in Campos Basin only in the
525
Miocene, being thus subsequent to the atypical hydrothermal petroleum migration. This
526
conventional petroleum system was responsible for the large volume of oil and gas that
527
fills the Pre-Salt reservoirs.
528
Salinities are consistently high in aqueous inclusions in the studied hydrothermal
529
phases (12 to 26 wt. % eq. NaCl; Table 1), suggesting that they were formed from
530
hypersaline brines. In this context, significant Th variation within a narrow salinity
531
range indicates temperature variation within the hydrothermal system during mineral
23 532
precipitation (Fig. 14). Nevertheless, no correlation was observed between Th and
533
salinity data. This, together with homogeneous high salinity values, indicates that there
534
was no fluid mixing in the studied hydrothermal system, and that the minerals
535
precipitated from a single hydrothermal system. Rare salinities below 15 wt. % eq.
536
NaCl from aqueous FIs hosted in macrocrystalline calcite and quartz (Table 1; Fig. 14)
537
could indicate some incipient isothermal mixing between the predominant hydrothermal
538
fluid (approximately 20–21 wt. % eq. NaCl) with a moderate-salinity fluid
539
(approximately 13 wt. % eq. NaCl). However, since 97% of the salinity results are
540
above 15 wt. % eq. NaCl, this relatively lower values may be related to salinity
541
variation in recharge area (Khaska et al., 2013).
542
The Th data in the range 92 to 152 °C with salinities of 13 to 26 wt. % eq. NaCl
543
of Campos Basin Pre-Salt are very similar to the majority of hydrothermal systems
544
observed in the hydrothermal dolomites (HTD), carbonate-hosted Pb-Zn Mississippi
545
Valley (MVT) and Irish-type deposits (Davies and Smith, 2006; Leach et al., 2001;
546
Wilkinson, 2001; Paradis et al., 2007; Wilkinson, 2010).
547
5.2. Carbon, oxygen and strontium isotopes
548
The δ13C values of host rock mineral phases (MBCB, MBCL, MBDD, FC and
549
CS) and hydrothermal mineral phases (SD and MC) present a very narrow variation
550
(between -1.88 and 2.43‰; Fig. 13, Table 2). These values are similar to those
551
previously reported from Pre-Salt carbonates from Campos, Santos and Kwanza Basins.
552
δ13C bulk-rock data from the cuttings of well CP-5 in Campos Basin vary between -0.38
553
and 2.71‰ for the Macabu Formation, and between -0.88 and 1.35‰ for the Coqueiros
554
Formation (Dias, 1998; Rodrigues, 2005; Muniz and Bosence, 2015). The δ13C bulk-
555
rock data (-1.52 to 3.20‰) of spherulite, shrub, laminate, calcarenite and calcirudite
24 556
lithofacies from core samples of the sag Barra Velha Formation from Santos Basin
557
(Farias et al., 2019) are similar to reported values of lithofacies from the Macabu
558
Formation (Herlinger Jr. et al., 2017; Lima and De Ros, 2019).
559
Syn-rift and sag intervals from Kwanza Basin have lithofacies and mineral
560
phases similar to those described for the rift (Coqueiros Formation) and sag (Macabu
561
Formation) intervals of the Campos Basin (Fig. 2). The δ13C values of spherulitic and
562
shrubby calcite and early dolomite from the sag interval of Kwanza Basin vary between
563
-4.3 and 3‰, whereas neomorphic calcite and calcite cement filling mollusk molds and
564
interparticle pores from syn-rift interval of Kwanza Basin usually display δ13C between
565
-3.0 and 2.0‰ (Saller et al., 2016; Sabato Ceraldi and Green, 2016; Fig. 13). Although
566
Sabato Ceraldi and Green (2016) present δ13C and δ18O values from rift and sag Pre-Salt
567
sections from Kwanza Basin, it is not possible to identify the carbonate analyzed. The
568
lithofacies, mineralogy and δ13C data of Campos, Santos and Kwanza basins suggest
569
that the depositional and diagenetic processes were quite similar within the vast Aptian
570
lacustrine system, indicating either a single water body or hydrogeological connected
571
lakes.
572
Data from the present study show that δ18O values of host rock mineral phases
573
(MBCB, MBCL, MBDD and CS) vary between -2.11 and 1.74‰, except for FC that
574
has values raging between -5.70 and -1.93‰ (Fig. 13, Table 2). Similar pattern was
575
obtained by Sabato Ceraldi and Green (2016). In contrast, hydrothermal mineral phases
576
(SD and MC) present much lower isotope values (-9.74 to -4.92‰; Fig. 13, Table 2),
577
suggesting either differences in the temperature of the fluid and/or in their oxygen
578
isotopic composition (δ18Ow). In addition, δ18O values of FC samples suggest
579
recrystallization with temperature and/or δ18Ow different from both other host rock
580
mineral phases and hydrothermal mineral phases. The comparison between the δ18O of
25 581
carbonates from the Pre-Salt and others formed elsewhere is not straightforward,
582
because their isotopic composition depends on the temperature, carbonate phase and
583
fluid δ18Ow (Sharp, 2007). Equations (1) and (2) show the relationship between δ18Ocarb,
584
δ18Ow and temperature for calcite (O’Neil et al., 1969) and dolomite (Matthews and
585
Katz, 1977), respectively:
586
10 ∝ = 2.78 ×
587
10 ∝ = 3.06 ×
− 2.89 (1) − 3.24
(2)
588
Where ( 10 ∝ ) is approximately equal to δ18Ocarb - δ18Ow, and T is the
589
temperature in Kelvin. These equations were used to calculate the δ18Ow of fluid in
590
equilibrium with each mineral phase (FC, CS, SD and MC) based on Th obtained from
591
aqueous and petroleum FIs, similar to the approach by King and Goldstein (2018). The
592
δ18Ow VPDB (Vienna Pee Dee Belemnite) was then converted to δ18Ow VSMOW
593
(Vienna Standard Mean Ocean Water) by using the equation 3 (O’Neil et al., 1969):
594
δ# O% &'()* = 1.03086 × δ# O% &+,- + 30.86 (3)
595
Calculated δ18Ow for FC samples vary between 5.4 and 12.4‰ (Table 4) based
596
on the maximum and minimum measured δ18Ocarb (-5.7 and -1.93‰) and Th derived
597
from petroleum FIs (83–111 °C). A variation of 6.5 to 10.4‰ was calculated for FC
598
δ18Ow using just two measurements of Th from aqueous FIs, both of 92.2 °C. Calculated
599
δ18Ow of CS samples vary between 6.3 and 16.3‰ based on the maximum and
600
minimum measured δ18Ocarb (-2.11 and 1.74‰) and Th derived from petroleum FIs (63–
601
112.2 °C). Those values calculated from Tho suggest that δ18Ow was the oxygen isotope
602
composition of pore water during the recrystallization of both mineral phases (FC and
603
CS). These high δ18Ow values are similar to those calculated for the isotopic
604
composition of fluids in equilibrium with shrub and spherulite lithofacies from Well
26 605
SB-2 of the Barra Velha Formation, Santos Basin (Farias et al., 2019). Based on oxygen
606
isotope compositions and clumped temperatures, the recalculated δ18Ow in equilibrium
607
with carbonates from these lithofacies vary between 5.0 to 10.7‰. These highly
608
enriched δ18Ow are comparable to the isotopic composition of residual evaporative
609
brines (Epstein and Mayeda, 1953) and/or of fluids resulted from extensive rock-water
610
interaction with basinal materials (King and Goldstein, 2018). Therefore, we suggest
611
that those values represent δ18Ow modified from residual evaporative brines during
612
burial, in which δ18Ow is the product of extensive rock-water interaction and
613
recrystallization. Considering this scenario, the FIs in spherulites and shrubs were
614
trapped during the recrystallization of these constituents during advanced burial. Some
615
FIs in these phases may, however, have been inherited from earlier stages of
616
neomorphism/recrystallization.
617
The oxygen isotopic composition of aqueous fluids (δ18Ow) in equilibrium with
618
MC are calculated between 3.3 and 11.1‰ (Table 4), based on the maximum and
619
minimum measured δ18Ocarb (-9.74 and -4.92‰) and on Th derived from aqueous FIs
620
(101−130 °C). These values are similar to those calculated for the aqueous fluids in
621
equilibrium with SD (δ18Ow between 2.3 and 10.0‰), which are based on aqueous FIs
622
in the range of 105 to 150 °C and maximum and minimum measured δ18Ocarb between -
623
9.49 and -6.35‰. These results indicate that MC and SD were formed from the same
624
hydrothermal fluid.
625
The
87
Sr/86Sr ratios from of host rock mineral phases (FC, CS and MBDD –
626
from 0.71299 to 0.71394) are significantly higher than those from hydrothermal mineral
627
phases (SB, SD and MC – from 0.71100 to 0.71226) (Fig. 13, Table 3). Reported
628
87
629
2017), which is the same well labeled as well A in this study, were obtained by leaching
Sr/86Sr values for the Macabu Formation (0.712991−0.713365) in well C (Tedeschi,
27 630
carbonate phases with weak acid, without differentiation of carbonate phases or micro-
631
sampling. The results presented here are also similar to previous data from other
632
Brazilian Pre-Salt wells at Campos and Santos basins (Dias, 1998; Pietzsch et al., 2018;
633
Farias et al., 2019; Fig. 13F).
634
The
87
Sr/86Sr ratios of carbonate samples from the rift Coqueiros and sag
635
Macabu formations range from 0.7117 to 0.7118 (n=2) and 0.7128 to 0.7137 (n=2)
636
(Dias, 1998). The
637
spherulite, shrub, laminate, calcarenite and calcirudite, from Santos Basin SB-2 well
638
vary from 0.7130 to 0.7138 (Farias et al., 2019). In addition, the 87Sr/86Sr ratios of 194
639
sidewall samples from well S10 in Santos Basin vary between 0.7105 and 0.7140
640
(Pietzsch et al., 2018). The isotope ratios from the rift Itapema Formation (Jiquiá local
641
stage) were usually lower than 0.7115, while the ratios from the sag Barra Velha
642
Formation (Alagoas local stage) were roughly between 0.7120 and 0.7140. These values
643
agree with five diagenetically modified ostracod samples analyzed by Tedeschi (2017),
644
three of which are from Hourcqia sp. valves (Itapema Formation;
645
0.711553 to 0.712012) and two are from Pattersoncypris ssp. valves (Barra Velha
646
Formation;
647
isotopes and 87Sr/86Sr data from different studies, depositional and diagenetic processes
648
were similar over a regional scale in Campos and Santos Basins, suggesting a single
649
vast lake or lakes that were hydrogeologically connected. Additionally, these data
650
establish constraints on
651
comparison with the hydrothermal mineral phases.
652
The
87
Sr/86Sr ratios from 43 samples from the sag phase, including
87
Sr/86Sr ratios from
87
Sr/86Sr ratios from 0.712756 to 0.713427). As indicated by the stable
87
87
Sr/86Sr ratios at a regional scale that can be used for
Sr/86Sr ratio of Pre-Salt carbonates indicate a strong interaction with 87
Sr/86Sr ratios,
653
continental crust materials, being significantly higher than marine
654
specifically than seawater values from Early Cretaceous (from 0.707074 to 0.707541;
28 655
Jenkyns et al., 1995; McArthur et al., 2012; Bodin et al., 2015; Yamamoto et al., 2013;
656
Ando, 2015) as well as that of Phanerozoic seawater (<0.7095; McArthur et al., 2012;
657
Fig. 13F). Based on Sr isotopes, Pietzsch et al. (2018) presented a geological model to
658
discriminate the Sr source of lacustrine rift and sag sections in the Santos Basin.
659
According to this model, Sr isotopes of carbonate rocks from sag Barra Velha
660
Formation were controlled by groundwater with long residence time of interaction with
661
felsic basement rocks. Although one may argue that leaching of waters through felsic
662
basement was not enough to deliver ions (Mg2+, Ca2+ and others) to develop the Barra
663
Velha Formation, a mass balance using waters from the basement has not been
664
performed. Therefore, although this process seems to be feasible at local scale, as
665
pointed out by the geochemical models of Teboul et al. (2016), the isotopic and
666
elemental aspects of the huge Pre-Salt volumes still need to be explained in basinal
667
scale.
668
Otherwise, coquinas from the Itapema Formation would have been deposited in
669
a rift lacustrine environment with higher fluvial recharge, and more variable catchment
670
area and subsurface recharge. This hydrological model for the Itapema Formation would
671
allow mixing of groundwater recharge with long residence time with waters derived
672
from Cabiúnas volcanic and volcaniclastic rocks and Piçarras Formation siliciclastic
673
sediments. Considering the 87Sr/86Sr ratios obtained from the Campos Basin, we would
674
consider the same hydrological patterns to explain the differences between
675
ratios from the rift Coqueiros and the sag Macabu formations.
676
The
87
Sr/86Sr
87
Sr/86Sr values of SD, MC and SB overlap each other (Fig. 13, Table 3).
677
Although the petrographic sequence shows minor dissolution of SD followed by
678
formation of MC and SB, there is not a definitive petrographic evidence that these
679
mineral phases are derived from different fluids (Figs. 5, 8). The same 87Sr/86Sr values
29 680
and the absence of petrographic evidence suggest that SD, MC and SB were formed
681
from the same fluid. In addition, strontium isotope composition of carbonate
682
hydrothermal phases (SD and MC) shows respectively positive and negative correlation
683
with δ13C and δ18O values, which also suggest a same trend of evolution from the same
684
fluid (Fig. 13). The relationship between δ13C and
685
observed in hydrothermal baroque dolomite from Cambrian-Ordovician Arbuckle
686
Group in USA midcontinent, whereas the relationship between δ18O and 87Sr/86Sr ratios
687
is the same (King and Goldstein, 2018). These observations suggest the same strontium
688
source for all these three mineral phases. The negative correlation between δ18O and
689
87
690
different rocks or that hydrothermal fluids at higher temperature and/or with lower
691
δ18Ow would promote stronger host rock dissolution, becoming more radiogenic.
692
Likewise, the positive correlation between
693
increased host rock dissolution by the hydrothermal fluids. According to this
694
interpretation, strong dissolution of basement and rift rocks precede the formation of
695
these hydrothermal mineral phases (SD, MC and SB).
696
87
Sr/86Sr is the opposite of those
Sr/86Sr (Fig. 13) suggests either that different set of fluids were in equilibrium with
87
Sr/86Sr and δ13C would be related to
5.3. Hydrothermal alteration of carbonate rocks
697
The term hydrothermal alteration was originally applied to situations with the
698
presence of hot waters and ore deposits associated with magmatic activities (e.g.,
699
Gilbert, 1875; Morey and Niggli, 1913; Holmes, 1928; Stearns et al., 1935). More
700
recently, hydrothermal alteration was defined as a replacement of the original minerals
701
by hydrothermal fluids that deliver reactants and remove aqueous reaction products
702
(Utada, 1980; Henley and Ellis, 1983; Inoue, 1995; Reed, 1997). Hydrothermal fluids
703
represent aqueous solutions of different compositions and origins (magmatic,
30 704
metamorphic, groundwater, or meteoric waters), which are hotter than the wall-rock
705
(Skinner, 1997).
706
Hydrothermal alterations represent the evidence of a geothermal anomaly,
707
requiring both a mechanism (heat source) and conduits for the flow of fluids (e.g. deep
708
fault systems). Hydrothermal systems are characterized by focusing of fluids through
709
fault systems at very high flow rates, in a transient or episodic regime (Davies, 2004).
710
Such focalization of fluids and their interaction with the host rocks are linked to the
711
origin of carbonate-hosted base metal deposits (e.g., the Pb-Zn Mississippi Valley type
712
deposits; Leach, 1994) and may also modify significantly the original properties of
713
some hydrocarbon reservoirs (e.g., Smith, 2006; Zhu et al., 2015). The so-called
714
hydrothermal petroleum is generated in volcanic and geothermal marine and continental
715
environments, and sampled from submarine and sublacustrine oil seeps (vent fluids),
716
mud volcanoes, hot springs, chimneys, fumaroles and other surface manifestations
717
(Ventura et al., 2012; Simoneit, 2018).
718
Despite several studies on hydrothermal ore deposits, the characterization and
719
definition of hydrothermal processes in sedimentary successions are still a matter of
720
heated discussion (e.g., Machel and Lonnee, 2002; Davies, 2004; Davies and Smith,
721
2006). White (1957) defined hydrothermal processes solely based on a temperature
722
contrast of 5 °C or more between the inflowing aqueous solutions and the wall-rocks,
723
disregarding any temperature limits or source of the fluids. Similarly, Machel and
724
Lonnee (2002) suggested that minerals formed in sedimentary rocks should be
725
considered hydrothermal minerals only if precipitated from fluids with temperatures at
726
least 5 to 10 °C hotter than the host rock temperature, disregarding the fluid sources and
727
the pathways. More loosely, Davies and Smith (2006) defined hydrothermal fluids as
728
those arising to the surface at temperatures higher than those of the depositional
31 729
environment, or that are introduced into the host rocks at a higher temperature than the
730
host. Following such line of definition, sedimentary ore-forming hydrothermal systems
731
may be either depositional (syngenetic) or diagenetic (epigenetic), depending on the
732
time, depth, geometry and mechanism of the flow of fluids. In still other lines of
733
definition, the minerals are hydrothermal if considered “unusual”, or “exotic” in relation
734
to the common carbonate rock mineralogy, including saddle dolomite, fluorite, barite,
735
anhydrite, dickite, sphalerite, and pyrite (Neilson and Oxtoby, 2008).
736
As reliable determination of paleotemperature contrasts in the order of only 5 to
737
10 °C is impossible in most cases, and as minerals that may be considered “exotic” by
738
some authors are widespread in sedimentary successions, or simply do not occur in
739
several recognized hydrothermal systems, we opted for a more practical and operational
740
definition for hydrothermal processes and products. This definition is based on the
741
focusing of the flow of relatively heated fluids coming from below through faults,
742
fractures, unconformities and similar conduits, and on the consequent concentration of
743
mineral precipitation and/or dissolution in the proximity of such conduits. Therefore, in
744
this sense, temperature and/or compositional requirements are not mandatory for the
745
definition of hydrothermal alterations.
746
5.3.1. Hydrothermal dolomitization
747
Hydrothermal dolomitization was defined by Davies and Smith (2006) as
748
occurring under subsurface conditions, commonly at shallow burial depths (< 1 km),
749
and developed along structural lineaments by hypersaline fluids presenting higher
750
temperatures and pressures than those of the host formation, which are usually
751
carbonate rocks. According to these authors, reservoir facies formed by structurally-
752
controlled hydrothermal dolomitization (HTD) are the largest hydrocarbon producers in
753
North America and have great potential in other basins. King and Goldstein (2018)
32 754
interpreted basin-derived hydrothermal system as responsible for baroque dolomite
755
precipitation with elevated Th and salinities in the Cambrian-Ordovician Arbuckle
756
Group. Hydrothermal calcite cement filling secondary porosity in hydrothermal
757
dolomites has also been described by some authors (e.g., Coveney Jr. et al. 2000,
758
Lonnee and Machel, 2006, Biehl et al., 2016, Mansurbeg et al., 2016).
759
Hydrothermal dolomitization has been reported in the Ordovician of the United
760
States, Canada, and China (e.g., Al-Aasm, 2003; Davies and Smith, 2006; Luczaj, 2006;
761
Luczaj et al., 2006; Conliffe et al., 2010; Xu et al., 2015), the Devonian and
762
Mississippian basins of Canada (e.g., Boreen and Colquhoun, 2001; Boreen and Davies,
763
2004), the Carboniferous of the United States and Spain (e.g., Gasparrini et al., 2006;
764
Hiemstra and Goldstein, 2015), the Permian of Germany (e.g., Biehl et al., 2016), the
765
Mesozoic of the Atlantic passive margins (e.g., Wierzbicki et al., 2006), the Jurassic-
766
Cretaceous of Spain, the Persian Gulf in Iraq (e.g., Mansurbeg et al., 2016), and Saudi
767
Arabia (e.g., Cantrell et al., 2004). Based on petrography, isotopic, fluid inclusions
768
characteristic, and on the operational definition previously proposed, the analyzed
769
saddle or baroque dolomites (Figs. 5, 6, 7, 8, 9, 11, 12, 13, 14) are defined as
770
hydrothermal, as also the extensive dolomitization observed in part of the associated
771
Pre-Salt rocks.
772
5.3.2.
Hydrothermal silicification
773
Silicification is a common diagenetic process, affecting a wide variety of
774
originally non-siliceous sediments (e.g., Namy, 1974; Meyers, 1977; Meyers and James,
775
1978; Hesse, 1989). Hydrothermal silicification can affect a wide variety of rock types
776
in multiple environments and geological contexts. However, few studies specifically
777
addressing the hydrothermal silicification of carbonate deposits have been published
778
(e.g., Bellanca et al., 1984; You et al., 2018). In most studies on the hydrothermal
33 779
alteration of carbonate sequences, silicification is considered a subordinate process in
780
comparison to dolomitization (e.g., Packard et al., 2001). Nevertheless, hydrothermal
781
silicification of Pre-Salt carbonates has been previously described in Campos Basin
782
(Vieira de Luca et al., 2017; Lima and De Ros, 2019), as well as of the African basins
783
(Poros et al., 2017; Teboul et al., 2017).
784
An example of the association of hydrothermal silicification and HTD
785
corresponds to the Upper Devonian (Famenian) Wabamun Group gas-condensate
786
reservoirs of the Parkland Field, Canada, which were generated by hydrothermal
787
dolomitization, silicification and dissolution (Packard et al., 2001). In those reservoirs,
788
hydrothermal silica occurs as lenses or as diffuse and discontinuous masses of
789
microquartz associated or not with dolomitization, as well as microquartz cement filling
790
fractures after saddle dolomite precipitation. The dolomitization and silicification of
791
those carbonate reservoirs occurred in relatively rapid succession and at shallow burial
792
depths. The pervasive precipitation of microcrystalline silica, macrocrystalline quartz
793
and spherulitic chalcedony observed in the study area (Figs. 5, 6, 7, 9, 11, 12, 14) was
794
defined as hydrothermal, based on the petrographic and fluid inclusions characteristics,
795
and on the operational definition previously presented.
796
5.4. Timing and source of the hydrothermal system
797
The hydrothermal processes recorded in northern Campos Basin Pre-Salt
798
lacustrine carbonates are imprinted in rocks of the sag section, and therefore cannot be
799
related to the rift magmatism. The scarcity of hydraulic breccias, the distribution of the
800
hydrothermal phases in the host carbonates, and the high homogenization temperature
801
(Th) of FIs indicate a hydrothermal system was active during effective burial, and not
802
related to the Aptian magmatism either. This indicates that the hydrothermal fluids
34 803
responsible for dissolution and for precipitation of mineral phases filling fractures that
804
cross-cut the studied carbonates are possibly related to Late Cretaceous and/or
805
Paleogene magmatic events.
806
The timing of the hydrothermal processes affecting northern Campos Basin Pre-
807
Salt deposits can be better evaluated within the framework of the burial and thermal
808
history of the area (Fig. 15). Based on this, and on petrographic and FIs evidence, the
809
observed hydrothermal alterations can be more probably related to the Late Cretaceous
810
(Santonian/Campanian), Paleocene and/or Eocene magmatic events (Fig. 15). The burial
811
and thermal history of the area indicates that during the advent of these magmatic
812
activities, the Pre-Salt reservoirs were at 2 to 4 kilometers below the seafloor and at 76
813
to 98 °C (Fig. 15). This temperature range is substantially lower (> 5–10 °C) than those
814
recorded in aqueous FIs hosted in the hydrothermal phases, thus indicating another heat
815
source to the system besides burial also had existed.
816
The δ13C data of hydrothermal SD and MC are similar (between -1.88 and
817
+1.28‰; Fig. 13, Table 2), suggesting both mineral phases were precipitated by the
818
same hydrothermal fluids, and from the same carbon source. Except for a few samples,
819
these hydrothermal phases are within the range of δ13C values of host rock mineral
820
phases (MBCB, MBCL, MBDD, FC and CS), which vary between 0.05‰ and 2.43‰.
821
Therefore, the carbon composition of the hydrothermal fluids was near to isotope
822
equilibrium with the host rock mineral phases. The slightly wider range of δ13C values
823
observed in SD samples could be explained by their wide temperature formation range
824
indicated by the FI data (105 to 150 °C, Table 1). For instance, the carbon isotopic
825
fractionation between calcite and CO2 vary from 1.0‰ at 150 °C to 3.4‰ at 100 °C
826
(Bottinga, 1969). This implies that temperature decrease alone could explain the carbon
35 827
isotopic range observed in SD, and that the isotopic composition of hydrothermal fluids
828
was buffered by their interaction with the host rocks.
829
Besides temperature, the observed variation in carbon isotopes could also be
830
related to mixture of fluids from different sources. The lowest δ13C value observed in
831
SD (-1.88‰) could suggest an original carbon isotopic for the hydrothermal fluids
832
before fluid-rock interaction, which is lower than 0.05‰. One possibility is that the
833
fluid had a δ13C value of -1.88‰. However, it is likely that this value is a result of fluid
834
with lower δ13C value and a higher δ13C value from the host-rock (mineral phases
835
usually display δ13C > 0.05‰). This interaction occurs by dissolving partially the
836
carbonates present in the host-rock. The dissolution of carbonate minerals from host
837
rock could be due to the low pH and/or low partial pressure of CO2 of the original fluid
838
before interaction with the host rock. As dissolution increases, higher is the CO2 from
839
host rock in modified hydrothermal fluid. This process is sensitive to temperature as the
840
system involves the interaction of H2O and CO2.
841
A possible original CO2 source for the hydrothermal fluids would be the
842
serpentinization of upper mantle (Fig. 16), occurring below areas with strongly thinned
843
continental crust or even through direct mantle exhumation (e.g., Boillot et al., 1987;
844
Manatschal and Bernoulli, 1999; Whitmarsh et al., 2001; Kusznir and Karner, 2007).
845
Mantle exhumation can liberate large volumes of Si, Ca, Mg and CO2 to the interacting
846
fluids (Frost and Beard, 2007; Pinto et al., 2015; 2017). Additionally, the exothermic
847
chemical reactions related to serpentinization (Moody, 1976; Proskurowski et al., 2006)
848
may generate hydrothermal fluids with high content of hydrogen and light hydrocarbons,
849
mainly CH4 (Allen and Seyfried Jr., 2004). Mantle exhumation would have occurred in
850
deep offshore southern and southeastern Brazilian basins, including Campos, during the
851
initial Atlantic Ocean opening, as suggested by seismic and gravimetric evidence
36 852
(Unternehr et al., 2010; Gomes et al., 2011; Zalán et al., 2011; Kumar et al., 2013;
853
Peron-Pinvidic et al., 2013; Kukla et al., 2018). Despite the lowest δ13C value observed
854
in SD (-1.88‰) is significantly higher than mantle isotopic composition, which shows a
855
peak carbon isotope signature of -5‰ (Deines, 2002), we cannot rule out this hypothesis.
856
Nevertheless, the 87Sr/86Sr ratios of both the syngenetic and the diagenetic/hydrothermal
857
Pre-Salt carbonates are much higher than the values expected for the derivation of Sr
858
(and of Ca, which has similar geochemical behavior; Banner, 1995) from mantle
859
exhumation and serpentinization.
860
More probable source for these δ13C values involves the contribution of light
861
carbon derived from organic matter alteration by the hydrothermal fluids, or their
862
interaction with sediments with lighter isotopic composition, such as the rift carbonates
863
from the Coqueiros Formation (e.g., well CP-5 of Dias, 1998; Fig. 16), and/or laminite
864
and spherulite lithofacies similar to those from Santos Basin Barra Velha Formation
865
(e.g., well SB-2 of Farias et al., 2019). An additional mechanism for the lowest δ13C
866
values could include the alteration of organic matter by thermal sulfate reduction (TSR;
867
Machel, 2001). The presence of sulfides and sulfates derived presumably from the same
868
hydrothermal fluid could suggest a possible role of TSR during hydrothermal fluid-rock
869
interaction. However, the carbon isotope values are not as negative as expected for this
870
process, such as in the hydrothermal calcite from Cambrian-Ordovician Arbulckle
871
Group (King and Goldstein, 2018).
872
Calculated δ18Ow from hydrothermal mineral phases partially overlap the values
873
calculated for fluids in equilibrium with the host rock mineral phases (FC and CS),
874
suggesting that part of the oxygen isotope variations can be explained by processes
875
involving dissolution of host rock during hydrothermal flow, and/or mixing between the
876
host pore waters and the hydrothermal fluids. The exceptions are the lowest calculated
37 877
δ18Ow values (roughly between 3 and 5‰), which may be related to fluids of a different
878
source or with quite distinct temperature. If the fluids had different sources, those values
879
enriched in δ18Ow would suggest extensive water-rock interaction along the fluid
880
migration pathways. Considering the geology of the basin, the Pre-Cambrian basement
881
(mainly granitic-gneissic felsic rocks), volcanic rocks from the Cabiúnas Formation
882
(mainly basaltic) and sediments from the Atafona and Coqueiros formations could have
883
interacted with the hydrothermal fluids (Fig. 16). The definition of the main rocks that
884
interacted with the hydrothermal fluids is however difficult to define based exclusively
885
on oxygen isotopes.
886
As hydrothermal fluids are usually a result of mixing between their ‘original’
887
composition and the host-rock fluids and minerals, it is likely that such fluid-rock
888
interaction would influence the
889
However, the
890
significantly higher than those from hydrothermal mineral phases (SD, MC and SB) and
891
do not overlap them. In addition, SD and SB mineral phases display high total strontium
892
content (Supplementary material). Thus, it is likely that strontium concentration of the
893
hydrothermal fluids were relatively higher than the host rock before the dissolution of
894
the latter and that SD 87Sr/86Sr ratios (~0.711) represent the ratio of hydrothermal fluids
895
before fluid-rock interaction. The
896
radiogenic than those from Cretaceous seawater, than fluids associated to hydrothermal
897
alteration of predominantly mafic igneous rocks worldwide (~0.7073; e.g., Burke et al.,
898
1982; Allègre et al., 2010), and than most of rift volcanic and volcaniclastic rocks from
899
Campos Basin Cabiúnas Formation (~0.708; Mizusaki et al., 1992; Tedeschi, 2017).
900 901
87
87
Sr/86Sr ratios of hydrothermal phases as well.
Sr/86Sr ratios of host rock mineral phases (FC, CS and BD) are
Assuming similar
87
87
Sr/86Sr ratios of SD are substantially more
Sr/86Sr ratios for the rift Coqueiros Formation to those of
Itapema Formation from Santos Basin (0.7105 to 0.7120; Tedeschi, 2017; Pietzsch et al.,
38 902
2018), such radiogenic values are likely to be related to the dissolution of the Coqueiros
903
Formation carbonates (Fig. 16). However, as the rift Itapema Formation has
904
ratios lower than those from the sag Barra Velha Formation (Pietzsch et al., 2018),
905
another source of fluids would be needed to explain such
906
syngenetic sag carbonates. The higher
907
diagenetic sag carbonates could represent a product of leaching of the Precambrian
908
felsic basement (Tupinambá et al., 2012; Teboul et al., 2016; Pietzsch et al., 2018). The
909
question is that the granitic-gneissic basement is not a suitable source for the huge
910
amounts of Mg and Ca precipitated in the sag rocks. Nevertheless, the lower
911
ratios observed in the analyzed hydrothermal phases in relation to the syngenetic and
912
early diagenetic sag carbonates could represent mixing with original fluids derived from
913
mantle serpentinization (Fig. 16), and/or a product of interaction with mafic volcanic
914
and volcanoclastic rocks from Cabiúnas Formation (Bertani and Carozzi, 1985a, b;
915
Misuzaki et al., 1992). Such mixing would be necessary to lower the 87Sr/86Sr ratios of
916
the hydrothermal mineral phases in comparison to the host rocks.
87
87
87
Sr/86Sr
Sr enrichment in the
Sr/86Sr ratios of the syngenetic and early
87
Sr/86Sr
917
According to the burial-thermal history (Fig. 15), the estimated maximum
918
temperature for the Pre-Salt reservoirs was approximately 110 °C during the occurrence
919
of the Eocene magmatic episode in the Campos Basin. This temperature is significantly
920
lower than the mean and maximum Th values, respectively 123.4 and 152 °C, obtained
921
in the aqueous FIs. The homogenization temperatures obtained in the hydrothermal
922
phases SD, MC, MQ and SB are even higher than the maximum burial temperatures
923
interpreted for the study area, including the present reservoir temperature. Therefore, the
924
Th data from SD, MC, MQ and SB represent further evidence of the hydrothermal
925
origin of these mineral phases, regardless their relation to specific magmatic periods.
926
Hydrocarbon FIs in these late-stage mineral phases suggest an atypical oil migration
39 927
(sensu Magoon and Dow, 1994) occurred in association with the hydrothermal fluid
928
flow. The absence of coexisting vapor- and liquid-rich FIs, diagnostic of phase-
929
separation effects (e.g., Goldstein and Reynolds, 1994; Jones et al., 1996; Moore et al.,
930
2001), indicates the lack of boiling processes within the studied Pre-Salt hydrothermal
931
system.
932
6. Conclusions
933
The overall geological and geochemical framework, combined with a specific
934
mineral paragenetic assemblage (saddle dolomite, macrocrystalline quartz, calcite, Sr-
935
barite, celestine, fluorite, dickite, sphalerite, galena, other metallic sulfides, and
936
bitumen) and FIs with corresponding range of salinities and homogenization
937
temperatures allowed us to recognize that the Pre-Salt carbonate reservoirs of the
938
northern Campos Basin were percolated by hydrothermal fluids chemically
939
(compositionally) comparable to those that formed in Mississippi Valley Type and
940
similar deposits.
941
The elevated homogenization temperatures (up to 152 °C) measured in
942
macrocrystalline calcite, mega-quartz, saddle dolomite and Sr-barite indicate
943
entrapment temperatures higher than the maximum values interpreted from the burial
944
history, confirming hydrothermal conditions during the late-stage alteration of the
945
studied reservoirs. The hydrothermal system characterized in northern Campos Basin
946
Pre-Salt reservoirs presents temperatures and salinities similar to Mississippi Valley and
947
Irish hydrothermal systems.
948
Saddle dolomite and macrocrystalline calcite precipitated from hydrothermal
949
fluids show lower δ18O values than the syngenetic and diagenetic carbonates. The
950
fascicular calcite aggregates, characteristic of Pre-Salt reservoirs, experienced strong
40 951
recrystallization under the action of the hydrothermal fluids and, therefore, present δ18O
952
values intermediate between the hydrothermal carbonates and the early diagenetic
953
spherulitic and microcrystalline calcite phases. All carbonates analyzed in the Pre-Salt
954
reservoirs of the northern Campos Basin present high 87Sr/86Sr, with lower ratios for the
955
hydrothermal saddle dolomite and macrocrystalline calcite than for the syngenetic,
956
diagenetic and recrystallized carbonates.
957
The high temperatures and salinities measured from the fluid inclusions, and the
958
isotopic data obtained from saddle dolomite, Sr-barite, macrocrystalline calcite and
959
mega-quartz indicate that the fault-focused hydrothermal system affecting the northern
960
Campos Basin Pre-Salt reservoirs probably involved mixing of fluids derived from
961
several sources, such as the interaction with the granitic-gneissic basement, the rift
962
sedimentary succession, the Late Cretaceous and Paleogene magmatism, and possibly
963
the rising and exhumation of the asthenosphere. The final hydrothermal fluid is a blend
964
in which end-members are not quantifiable.
965
The fluid inclusions, isotopic and petrographic data, such as the scarcity of
966
hydraulic breccia, integrated to the burial and thermal history of the study area are
967
strong evidence that the hydrothermal alteration of northern Campos Basin Pre-Salt
968
reservoirs occurred under relatively deep burial situation of more than 2 kilometers. The
969
studied hydrothermal alterations had strong impact on the porosity, permeability, and
970
heterogeneity, contributing, together with the associated fracturing, to the production
971
performance of the Pre-Salt reservoirs.
972
Acknowledgements
973
The results and interpretations of this paper are part of the PhD research project
974
of BEML, funded by Petróleo Brasileiro S.A. - Petrobras. The authors wish to thank
41 975
Petrobras, for supporting this study and for the opportunity to publish this paper. In
976
particular, we are extremely grateful to Gustavo Garcia, who provided useful insight
977
and discussions about the burial-thermal history. We are grateful for the analytical and
978
technical assistance provided by Petrobras Research Center (CENPES), the Federal
979
University of Goiás (UFG) Regional Center for Technological Development and
980
Innovation (CRTI), the University of Brasília (UnB) Geochronology Laboratory, and
981
the University of São Paulo (USP). We would also like to thank the support of the
982
Graduate Geosciences Program of Rio Grande do Sul Federal University (UFRGS). The
983
authors wish also to thank the reviewers for their constructive suggestions that helped
984
improve the manuscript.
985
References
986
Al-Aasm, 2003. Origin and characterization of hydrothermal dolomite in the Western
987
Canada Sedimentary Basin. Journal of Geochemical Exploration 78–79, 9–15.
988
Allègre, C.J., Louvat, P., Gaillardet, J., Meynadier, L., Rad, S., Capmas, F., 2010. The
989
fundamental role of island arc weathering in the oceanic Sr isotope budget. Earth and
990
Planetary Science Letters 292 (1–2), 51–56.
991
Allen, D.G., Seyfried Jr., W.E, 2004. Serpentinization and heat generation: constraints
992
from Lost City and Rainbow hydrothermal systems. Geochimica et Cosmochimica
993
Acta 68 (6), 1347–1354.
994
Alvarenga, R.S., Iacopini, D., Kuchle, J., Scherer, C.M.S., Goldberg, K., 2016. Seismic
995
characteristics and distribution of hydrothermal vent complexes in the Cretaceous
996
offshore rift section of the Campos Basin, offshore Brazil. Marine and Petroleum
997
Geology 74, 12–25.
42 998 999
Anderson, G.M., 1991. Organic maturation and ore precipitation in Southeast Missouri. Economic Geology 86 (5), 909–926.
1000
Ando, A., 2015. Intersite discrepancy in the amplitude of marine negative δ13C
1001
excursion at the onset of early Aptian oceanic anoxic event 1a: Reconciliation
1002
through Sr isotopic screening of peculiar diagenetic overprint on the Pacific
1003
reference section (Deep Sea Drilling Project Site 463). In: Neal, C.R., Sager, W.W.,
1004
Sano, T., Erba, E. (Eds.), Environmental consequences of Ontong Java Plateau and
1005
Kerguelen Plateau volcanism. Geological Society of America Special Paper 511. The
1006
Geological Society of America, Boulder, USA, pp. 329–339.
1007
Armelenti, G., Goldberg, K., Kuchle J., De Ros, L.F., 2016. Deposition, diagenesis and
1008
reservoir potential of non-carbonate sedimentary rocks from the rift section of
1009
Campos Basin, Brazil. Petroleum Geoscience 22 (3), 223–239.
1010 1011
Banner, J.L., 1995. Application of the trace element and isotope geochemistry of strontium to studies of carbonate diagenesis. Sedimentology 42 (5), 805–824.
1012
Baumgarten, C.S., Dutra, A.J.C., Scuta, M.S., Figueiredo, M.V.L., Sequeira, M.F.P.B.,
1013
1988. Coquinas da Formação Lagoa Feia, Bacia de Campos: evolução da geologia de
1014
desenvolvimento. Boletim de Geociências da Petrobras 2 (1), 27–36.
1015
Bellanca, A., Censi, P., Di Salvo, P., Neri, R., 1984. Textural, Chemical and Isotopic
1016
Variations Induced by Hydrothermal Fluids on Mesozoic Limestones in
1017
Northwestern Sicily. Mineralium Deposita 19, 78–85.
1018
Bertani, R.T., Carozzi, A.V., 1985a. Lagoa Feia Formation (Lower Cretaceous),
1019
Campos Basin, offshore Brazil: rift valley stage lacustrine carbonate reservoirs, I.
1020
Journal of Petroleum Geology 8, 37–58.
43 1021
Bertani, R.T., Carozzi, A.V., 1985b. Lagoa Feia Formation (Lower Cretaceous) Campos
1022
Basin, offshore Brazil: rift valley type lacustrine carbonate reservoirs, II. Journal of
1023
Petroleum Geology 8, 199–220.
1024
Bertrand, R., Chagnon, A., Malo, M., Duchaine, Y., Lavoie, D., Savard, M.M., 2003.
1025
Sedimentologic, diagenetic and tectonic evolution of the Saint-Flavien gas reservoir
1026
at the structural front of the Quebec Appalachians. Bulletin of Canadian Petroleum
1027
Geology 51 (2), 126–154.
1028
Biehl, B.C., Reuning, L., Schoenherr, J., Lüders, V., Kukla, P.A., 2016. Impacts of
1029
hydrothermal dolomitization and thermochemical sulfate reduction on secondary
1030
porosity creation in deeply buried carbonates: a case study from the Lower Saxony
1031
Basin, northwest Germany. American Association of Petroleum Geologists Bulletin
1032
100 (4), 597–621.
1033
Bodin, S., Meissner, P., Janssen, N.M.M., Steuber, T., Mutterlose, J., 2015. Large
1034
igneous provinces and organic carbon burial: Controls on global temperature and
1035
continental weathering during the Early Cretaceous. Global and Planetary Change
1036
133, 238–253.
1037 1038
Bodnar, R.J., 2003. Reequilibration of fluid inclusions. Fluid inclusions: Analysis and Interpretation, pp. 213–230.
1039
Bodnar, R.J., Reynolds, T.J., Kuehn, C.A., 1985. Fluid-inclusion systematics in
1040
epithermal systems. In: Berger, B., Bethke, R. (Eds.), Geology and Geochemistry of
1041
Epithermal Systems. Society of Economic Geologists, Reviews in Economic
1042
Geology 2, pp. 73–97.
1043
Bodnar, R.J., Vityk, M.O., 1994. Interpretation of microthermometric data for H2O-
1044
NaCl fluid inclusions. In: De Vivo, B., Frezzotti, M.L. (Eds.), Fluid Inclusions in
1045
Minerals, Methods and Applications. Virginia Tech Blackburg, pp. 117–130.
44 1046
Boillot, G., Recq, M., Winterer, E.L., Meyer, A.W., Applegate, J., Baltuck, M., Bergen,
1047
J.A., Comas, M.C., Davies, T.A., Dunham, K., Evans, C.A., Girardeau, J., Goldberg,
1048
G., Haggerty, J., Jansa, L.F., Johnson, J.A., Kasahara, J., Loreau, J.P., Luna-Sierra,
1049
E., Moullade, M., Ogg, J., Sarti, M., Thurow, J.,Williamson, M., 1987. Tectonic
1050
denudation of the upper mantle along passive margins: a model based on drilling
1051
results (ODP leg 103, western Galicia margin, Spain). Tectonophysics 132, 335–342.
1052
Boreen, T., Colquhoun, K., 2001. Ladyfern, NEBC: major gas discovery in the
1053
Devonian Slave Point Formation. Canadian Society of Petroleum Geologists, Annual
1054
Convention, Abstracts, 112–115.
1055
Boreen, T., Davies, G.R., 2004. Hydrothermal dolomite and leached limestones in a
1056
TCF gas play: the Ladyfern Slave Point reservoir, NEBC. In: McAuley, R. (Ed.),
1057
Dolomites - the spectrum: mechanisms, models, reservoir development. Canadian
1058
Society of Petroleum Geologists, Seminar and Core Conference, June 13–15,
1059
Calgary, Alberta, p. 17.
1060
Bottinga, Y., 1969. Calculated fractionation factors for carbon and hydrogen isotope
1061
exchange in the system calcite-carbon dioxide-graphite-methane-hydrogen-water
1062
vapor. Geochimica et Cosmochimica Acta 33, 49–64.
1063
Burke, W.H., Denison, R.E., Hetherington, E.A., Koepnick, R.B., Nelson, H.F., Otto,
1064
J.B., 1982. Variation of seawater 87Sr/86Sr throughout Phanerozoic time. Geology 10
1065
(10), 516–519.
1066
Callot, J.P., Breesch, L., Guilhaumou, N., Roure, F., Swennen, R., Vilasi, N., 2010.
1067
Paleofluids characterization and fluid flow modelling along a regional transect in
1068
Northern United Arab Emirates (UAE). Arabian Journal of Geosciences 3, 413–437.
1069 1070
Cantrell, D.L., Swart, P.K., Hagerty, R.M., 2004. Genesis and characterization of dolomite, Arab-D reservoir, Ghawar field, Saudi Arabia. GeoArabia 9 (2), 1–26.
45 1071
Castro, J.C., 2006. Evolução dos conhecimentos sobre as coquinhas-reservatório da
1072
Formação Lagoa Feia no trend Badejo-Linguado-Pampo, Bacia de Campos. Boletim
1073
de Geociências da Petrobras 25 (2), 175–186.
1074
Conliffe, J., Azmy, K., Gleeson, S.A., Lavoie, D., 2010. Fluids associated with
1075
hydrothermal dolomitization in St. George Group, western Newfoundland, Canada.
1076
Geofluids 10, 422–437.
1077
Coplen, T.B., Brand, W.A., Gehre, M., Gröning, M., Meijer, H.A.J., Toman, B.,
1078
Verkouteren, R.M., 2006. New guidelines for δ13C measurements. Analytical
1079
Chemistry 78 (7), 2439–2441.
1080
Corbella, M., Ayora, C., Cardellach, E., Soler, A., 2006. Reactive transport modeling
1081
and hydrothermal karst genesis: the example of the Rocabruna barite deposit (Eastern
1082
Pyrenees). Chemical Geology 233, 113–125.
1083
Coveney Jr., R.M., Ragan, V.M., Brannon, J.C., 2000. Temporal benchmarks for
1084
modeling Phanerozoic flow of basinal brines and hydrocarbons in the southern
1085
Midcontinent based on radiometrically dated calcite. Geology 28 (9), 795–798.
1086
Davies, G.R., 2002. Thermobaric dolomitization: Transient pressure driven processes
1087
and the role of boiling-effervescence in fabric/reservoir development (abs.).
1088
Canadian Society of Petroleum Geologists, 75th Anniversary Convention Abstracts
1089
Vol., p. 105.
1090
Davies, G.R., 2004. Hydrothermal (thermobaric) dolomitization: rock fabrics and
1091
organic petrology. In: McAuley, R. (Ed.), Dolomites - the spectrum: mechanisms,
1092
models, reservoir development. Canadian Society of Petroleum Geologists, Seminar
1093
and Core Conference, June 13–15, Calgary, Alberta, p. 17.
46 1094
Davies, G.R., Smith, L.B.J., 2006. Structurally controlled hydrothermal dolomite
1095
reservoir facies: an overview. American Association of Petroleum Geologists
1096
Bulletin 90 (11), 1641–1690.
1097 1098
Deines, P. 2002. The carbono isotope geochemistry of mantle xenoliths. Earth-Science Reviews 58 (1-4), 247–278.
1099
Dias, J.L., 1998. Análise sedimentológica e estratigráfica do Andar Aptiano em parte da
1100
Margem Leste do Brasil e no Platô das Malvinas: considerações sobre as primeiras
1101
incursões e ingressões marinhas do Oceano Atlântico Sul Meridional. (PhD Thesis).
1102
Universidade Federal do Rio Grande do Sul, Porto Alegre, Brazil, p. 208 (in
1103
Portuguese).
1104
Dias, J.L., Oliveira, J., Vieira, J., 1988. Sedimentological and stratigraphic analysis of
1105
the Lagoa Feia Formation, rift phase of Campos Basin, offshore Brazil. Revista
1106
Brasileira de Geociências 18, 252–260.
1107
Disnar, J.R., 1996. A comparison of mineralization histories for two MVT deposits,
1108
Treves and Les Malines (Causses Basin, France), based on the geochemistry of
1109
associated organic matter. Ore Geology Reviews 11 (1-3) 133–156.
1110 1111
Epstein, S., Mayeda, T., 1953. Variation of 18O content of waters from natural sources. Geochimica et Cosmochimica Acta, 4, 213–224.
1112
Farias, F., Szatmari, P., Bahniuk, A., França, A.B., 2019. Evaporitic carbonates in the
1113
pre-salt of Santos Basin – Genesis and tectonic implications. Marine and Petroleum
1114
Geology 105, 251–272
1115
Fontana, S., Nader, F.H., Morad, S., Ceriani, A., Al-Aasm, I.S., 2010. Diagenesis of the
1116
Khuff Formation (Permian-Triassic), northern United Arab Emirates. Arabian
1117
Journal of Geosciences 3 (4), 351–368.
47 1118 1119
Frost, B.R., Beard, J.S., 2007. On silica activity and serpentinization. Journal of Petrology, 48 (7), 1351–1368.
1120
Gasparrini, M., Bechstaedt, T., Boni, M., 2006. Massive hydrothermal dolomitization in
1121
the Southwestern Cantabrian Zone (Spain) and its relation to the late Variscan
1122
evolution. Marine Petroleum Geology 23, 543–568.
1123
Gilbert, G.K., 1875. Report on the geology of portions of Nevada, Utah, California, and
1124
Arizona. United States Geographic and Geological Surveys W. 100th Meridian 3,
1125
17–187.
1126
Girard, J.P., San Miguel, G., 2017. Evidence of High Temperature Hydrothermal
1127
Regimes in the Pre-Salt Series, Kwanza Basin, Offshore Angola. American
1128
Association of Petroleum Geologists Annual Convention and Exhibition. Houston,
1129
Texas, USA, Abstracts.
1130 1131
Goldstein, R.H., 2001. Fluid inclusions in sedimentary and diagenetic systems. Lithos 55, 159–193.
1132
Goldstein, R.H., Reynolds, T.J., 1994. Fluid Inclusion Microthermometry. In: Goldstein,
1133
R.H., Reynolds, T.J. (Eds.), Systematics of Fluid Inclusions in Diagenetic Minerals.
1134
Society for Sedimentary Geology, SEPM Short Courses 31, Tulsa, OK, USA, pp.
1135
87–121.
1136
Gomes, P.O., Kilsdonk, W., Grow, T., Minken, J., Barragan, R., 2011. Tectonic
1137
evolution of the Outer High of Santos Basin, southern São Paulo Plateau, Brazil, and
1138
implications for hydrocarbon exploration. In: Gao, D. (Ed.), Tectonics and
1139
sedimentation: Implications for petroleum systems. American Association of
1140
Petroleum Geologists, USA, Memoir 100, pp. 1–14.
48 1141
Greenwood, P.F., Brocks, J.J., Grice, K., Schwark, L., Jaraula, C.M.B., Dick, J.M.,
1142
Evans, K.A., 2013. Organic geochemistry and mineralogy. I. Characterization of
1143
organic matter associated with metal deposits. Ore Geology Reviews 50, 1–27.
1144 1145 1146 1147 1148 1149
Gregg, J.M., Sibley, D.F., 1984. Epigenetic dolomitization and the origin of xenotopic dolomite texture. Journal of Sedimentary Petrology 54, 908–931. Gregg, J.M., Sibley, D.F., 1987. Classification of dolomite rock textures. Journal of Sedimentary Petrology 57 (6), 967–975. Henley, R.W., Ellis, A.J., 1983. Geothermal systems ancient and modern: a geochemical review. Earth-Science Reviews 19, 1–50.
1150
Herlinger Jr., R., Zambonato, E.E., De Ros, L.F., 2017. Influence of diagenesis on the
1151
quality of lower cretaceous Pre-Salt lacustrine carbonate reservoirs from northern
1152
Campos Basin, offshore Brazil. Journal of Sedimentary Research 87, 1285–1313.
1153
Hesse, R., 1989. Silica diagenesis: origin of inorganic and replacement cherts. Earth-
1154
Science Reviews 26, 253–284.
1155
Hiemstra, E.J., Goldstein, R.H., 2015. Repeated injection of hydrothermal fluids into
1156
downdip carbonates: a diagenetic and stratigraphic mechanism for localization of
1157
reservoir porosity, Indian Basin Field, New Mexico, USA. In: Agar, S.M., Geiger, S.
1158
(Eds.), Fundamental controls on fluid flow in carbonates: current workflows to
1159
emerging technologies. Geological Society of London, Special Publications 406, pp.
1160
141–177.
1161
Holmes, A., 1928. The Nomenclature of Petrology. Ed. Van Nostrand-Reinhold, 284 p.
1162
Ilchik R.P., Brimhall G.H., Schull H.W., 1986. Hydrothermal maturation of indigenous
1163
organic matter at the Alligator Ridge Gold Deposits, Nevada. Economic Geology 81
1164
(1), 113–130.
49 1165 1166
Inoue, A., 1995. Formation of clay minerals in hydrothermal environments. In: Velde, B. (Ed.), Origin and Mineralogy of Clays. Springer-Verlag, Berlin, pp. 268−330.
1167
Jenkyns, H.C., Paull, C.K., Cummins, D.I., Fullagar, P.D., 1995. Strontium isotope
1168
stratigraphy of Lower Cretaceous atoll carbonates in the Mid-Pacific Mountains. In:
1169
Winterer, E.L., Sager, W.W., Firth, J.V., Sinton, J.M. (Eds.), Proceedings of the
1170
Ocean Drilling Program, Scientific Results 143. Texas A & M University, Ocean
1171
Drilling Program, College Station, USA, pp. 89–97.
1172
Jones, B., Renaut, R.W., Rosen, M.R., 1996. High-temperature (>90°C) calcite
1173
precipitation at Waikite Hot Springs, North Island, New Zealand. Journal of the
1174
Geological Society 153, 481–496.
1175
Khaska, M., Le Gal La Salle, C., Lancelot, J., team, ASTER, Mohamad, A., Verdoux,
1176
P., Noret, A., Simler, R., 2013. Origin of groundwater salinity (current seawater vs.
1177
saline deep water) in a coastal karst aquifer based on Sr and Cl isotopes. Case study
1178
of the La Clape massif (southern France). Applied Geochemistry 37, 212–227.
1179
King, B.D., Goldstein, R.H., 2018. History of hydrothermal fluid flow in the
1180
midcontinent,
1181
unconformities and porosity distribution. In: Armitage, P.J., Butcher, A.R., Churchill,
1182
J.M., Csoma, A.E., Hollis, C., Lander, R.H., Omma, J.E., Worden, R.H. (Eds.),
1183
Reservoir Quality of Clastic and Carbonate Rocks: Analysis, Modelling and
1184
Prediction. Geological Society of London, Special Publications 435, pp. 283–320.
1185
Kukla, P.A., Strozyk, F., Mohriak, W.U., 2018. South Atlantic salt basins – Witnesses
1186
USA:
the
relationship
between
inverted
thermal
structure,
of complex passive margin evolution. Gondwana Research 53, 41−57.
1187
Kumar, N., Danforth, A., Nuttall, P., Helwig, J., Bird, D.E., Venkatraman, S., 2013.
1188
From oceanic crust to exhumed mantle: a 40 year (1970–2010) perspective on the
1189
nature of crust under the Santos Basin, SE Brazil. In: Mohriak, W.U., Danforth, A.,
50 1190
Post, P.J., Brown, D.E., Tari, G.C., Nemcok, M., Sinha, S.T. (Eds.), Conjugate
1191
Divergent Margins. Geological Society of London, Special Publications 369, pp.
1192
147–165.
1193
Kusznir, N.J., Karner, G.D., 2007. Continental lithospheric thinning and breakup in
1194
response to upwelling divergent mantle flow: application to the Woodlark,
1195
Newfoundland and Iberia margins. In: Karner, G.D., Manatschal, G., Pinheiro, L.M.
1196
(Eds.), Imaging, Mapping and Modelling Continental Lithosphere Extension and
1197
Breakup. Geological Society of London, Special Publications 282, pp. 389–419.
1198
Leach, D.L., 1994. Genesis of the Ozark Mississippi Valley-type metallogenic province.
1199
In: Fontboté, L., Boni, M., (Eds.), Sediment Hosted Zn-Pb Ores. Springer-Verlag, pp.
1200
104–138.
1201
Leach, D.L., Bradley, D., Lewchuk, M.T., Symons, D.T.A., De Marsily, G., Brannon, J.,
1202
2001. Mississippi Valley-type lead-zinc deposits through geological time:
1203
Implications from recent age-dating research. Mineralium Deposita 36, 711–740.
1204
Lepley, S., Piccoli, L., Chitale, V., Kelley, I., Quest, M., 2017. The Importance of
1205
Understanding Diagenesis for the Development of Pre-Salt Lacustrine Carbonates.
1206
American Association of Petroleum Geologists Annual Convention and Exhibition.
1207
Houston, Texas, USA, Abstracts.
1208 1209
Leyden, R., Asmus, H., Zembruscki, S., Bryan, G., 1976. South Atlantic diapiric structures. American Association of Petroleum Geologists Bulletin 60 (2), 196–212.
1210
Lima, B.E.M., De Ros, L.F., 2019. Deposition, diagenetic and hydrothermal processes
1211
in Aptian Pre-Salt lacustrine carbonate reservoirs of the northern Campos Basin,
1212
offshore Brazil. Sedimentary Geology 383, 55–81.
1213
Lonnee, J., Machel, H.G., 2006. Pervasive dolomitization with subsequent hydrothermal
1214
alteration in the Clarke Lake gas field, Middle Devonian Slave Point Formation,
51 1215
British Columbia, Canada. American Association of Petroleum Geologists Bulletin
1216
90, 1739–1761.
1217
Luczaj, J.A., 2006. Evidence against the Dorag (mixing-zone) model for dolomitization
1218
along the Wisconsin arch – A case for hydrothermal diagenesis. American
1219
Association of Petroleum Geologists Bulletin 90 (11), 1719–1738.
1220
Luczaj, J.A., Harrison, W.B., Williams, N.S., 2006. Fractured hydrothermal dolomite
1221
reservoirs in the Devonian Dundee Formation of the central Michigan Basin.
1222
American Association of Petroleum Geologists Bulletin 90 (11), 1787–1801.
1223 1224
Machel, H.G., 2001. Bacterial and thermochemical sulfate reduction in diagenetic settings - old and new insights. Sedimentary Geology 140, 143–175.
1225
Machel, H.G., 2004. Concepts and models of dolomitization: a critical reappraisal. In:
1226
Braithwaite, C.J.R., Rizzi, G., Darke, G. (Eds.), The geometry and petrogenesis of
1227
dolomite hydrocarbon reservoirs, vol. 235, pp. 7–63. Geological Society of London,
1228
Special Publications.
1229 1230
Machel, H.G., Lonnee, J., 2002. Hydrothermal dolomite: a product of poor definition and imagination. Sedimentary Geology 152, 163–171.
1231
Machel, H.G., Mountjoy, E.W., 1987. General constraints on extensive pervasive
1232
dolomitization - and their application to the Devonian carbonates of western Canada.
1233
Bulletin of Canadian Petroleum Geology 35 (2), 143–158.
1234
Magoon, L.B., Dow, W.G., 1994. The Petroleum System. In: Magoon, L.B, Dow, W.G.
1235
(Eds.), Chapter 1: The petroleum system - from source to trap. American Association
1236
of Petroleum Geologists, USA, Memoir 60, pp. 3–24.
1237 1238
Manatschal, G., Bernoulli, D., 1999. Architecture and tectonic evolution of nonvolcanic margins: present day Galicia and ancient Adria. Tectonics 18, 1099–1119.
52 1239
Mansurbeg, H., Morad, D., Othmana, R., Morad, S., Ceriani, A., Al-aasm, I., Kolo, K.,
1240
Spirov, P., Proust, J.N., Preat, A., Koyi, H., 2016. Hydrothermal dolomitization of
1241
the Bekhme formation (Upper Cretaceous), Zagros Basin, Kurdistan Region of Iraq:
1242
Record of oil migration and degradation. Sedimentary Geology 341, 147–162.
1243 1244
Mattews, A.K., Katz, A., 1977. Oxygen isotope fractionation during the dolomitization of calcium carbonate. Geochimica et Cosmochimica Acta 41, 1431–1438.
1245
McArthur, J.M., Howarth, R.J., Shields, G.A., 2012. Chapter 7 – Strontium Isotope
1246
Stratigraphy. In: Gradstein, F.M., Ogg, J.G., Schmitz, M.D., Ogg, G.M. (Eds.), The
1247
Geologic Time Scale. Elsevier, Boston, USA, pp. 127–144.
1248
Mello, M.R., Mohriak, W.U., Koutsoukos, E.A.M., Bacoccoli, G., 1994. Selected
1249
petroleum systems in Brazil. In: Magoon, L.B., Dow, W.G. (Eds.), The petroleum
1250
system - from source to trap. American Association of Petroleum Geologists, USA,
1251
Memoir 60, pp. 499–512.
1252 1253 1254
Meyers, W.J., 1977. Chertification in the Mississippian Lake Valley Formation, Sacramento Mountains, New Mexico. Sedimentology 24 (1), 75–105. Meyers, W.J., James, A.T., 1978. Stable isotopes of cherts and carbonate cements in the
1255
Lake
Valley
Formation
1256
Sedimentology 25, 105–124.
(Mississipian),
Sacramento
Mts.,
New
Mexico.
1257
Mizusaki, A.M.P., Petrini, R., Bellieni, P., Comin-Chiaramonti, P., Dias, J.L., De Min,
1258
A., Piccirillo, E.M., 1992. Basalt magmatism along the passive continental margin of
1259
SE Brazil (Campos Basin). Contributions to Mineralogy and Petrology 111 (2), 143–
1260
160.
1261
Moody, J.B., 1976. Serpentinization: a review. Lithos 9 (2), 125–138.
53 1262
Moore, J.N., Norman, D.I., Kennedy, B.M., 2001. Fluid inclusion gas compositions
1263
from an active magmatic–hydrothermal system: a case study of The Geysers
1264
geothermal field, USA. Chemical Geology 173, 3–30.
1265 1266
Morey, G.W., Niggli, P., 1913. The hydrothermal formation of silicates, a review. Journal of the American Chemical Society 35, 1086–1130.
1267
Muniz, M.C., Bosence, D.W.J., 2015. Pre-Salt microbialites from the Campos Basin
1268
(offshore Brazil): image log facies, facies model and cyclicity in lacustrine
1269
carbonates. In: Bosence, D.W.J., Gibbons, K.A., LeHeron, D.P., Morgan, W.A.,
1270
Pritchard, T., Vining, B.A. (Eds.), Microbial carbonates in space and time:
1271
implications for global exploration and production. Geological Society, London,
1272
Special Publications 418, pp. 221–242.
1273 1274
Namy, J.N., 1974. Early diagenetic cherts in the Marble Falls Group (Pennsylvanian) of central Texas. Journal of Sedimentary Petrology 44, 1262–1268.
1275
Neilson, J.E., Oxtoby, N.H., 2008. The relationship between petroleum, exotic cements
1276
and reservoir quality in carbonates - a review. Marine Petroleum Geology 25, 778–
1277
790.
1278 1279
O’Neil, J.R., Clayton, R.N., Mayeda, T.K., 1969. Oxygen isotope fractionation in divalent metal carbonates. Journal of Chemical Physics 51, 5547–5558.
1280
Ostendorf, J., Henjes-Kunst, F., Mondillo, N., Boni, M., Schneider, J., Gutzmer, J.,
1281
2015. Formation of Mississippi Valley–type deposits linked to hydrocarbon
1282
generation in extensional tectonic settings: Evidence from the Jabali Zn-Pb-(Ag)
1283
deposit (Yemen). Geology 43 (12), 1055–1058.
1284
Packard, J.J., Al-Aasm, I.S., Samson, I., Berger, Z., Davies, J., 2001. A Devonian
1285
hydrothermal chert reservoir: the 225 bcf Parkland field, British Columbia, Canada.
1286
American Association of Petroleum Geologists Bulletin 85 (1), 51–84.
54 1287
Paradis, S., Hannigan, P., Dewing, K., 2007. Mississippi Valley-type lead-zinc deposits.
1288
In: Goodfellow, W.D. (Ed.), Mineral Deposits of Canada: A Synthesis of Major
1289
Deposit-Types, District Metallogeny, the Evolution of Geological Provinces, and
1290
Exploration Methods. Geological Association of Canada, Mineral Deposits Division,
1291
Special Publication 5, pp. 185–203.
1292
Parnell, J., Baron, M., Mann, P., Carey, P., 2003. Oil migration and bitumen formation
1293
in a hydrothermal system, Cuba. Journal of Geochemical Exploration 78-79, 409–
1294
415.
1295
Peron-Pinvidic, G., Manatschal, G., Osmundsen, P.T., 2013. Structural comparison of
1296
archetypal Atlantic rifted margins: A review of observations and concepts. Marine
1297
and Petroleum Geology 43, 21–47.
1298
Pietzsch, R., Oliveira, D.M., Tedeschi, L.R., Queiroz Neto, J.V., Figueiredo, M.F.,
1299
Vazquez, J.C., Souza, R.S., 2018. Palaeohydrology of the Lower Cretaceous pre-salt
1300
lacustrine system, from rift to post-rift phase, Santos Basin, Brazil. Palaeogeography,
1301
Palaeoclimatology, Palaecology 507, 60–80.
1302
Pinto, V.H.G., Manatschal, G., Karpoff, A.M., Viana, A., 2015. Tracing mantle-reacted
1303
fluids in magma-poor rifted margins: The example of Alpine Tethyan rifted margins.
1304
Geochemistry Geophysics Geosystems 16, 3271–3308.
1305
Pinto, V.H.G., Manatschal, G., Karpoff, A.M., Ulrich, M., Viana, A.R., 2017. Seawater
1306
storage and element transfer associated with mantle serpentinization in magma-poor
1307
rifted margins: a quantitative approach. Earth and Planetary Science Letters 459,
1308
227–237.
1309
Poros, Z., Jagniecki, E., Luczaj, J., Kenter, J., Gal, B., Correa, T.S., Ferreira, E.,
1310
McFadden, K.A., Elifritz, A., Heumann, M., Johnston, M., Matt, V., 2017. Origin of
1311
silica in Pre-Salt carbonates, Kwanza Basin, Angola. AAPG Annual Convention and
55 1312
Exhibition, Houston, Texas, April 2–5. USA: AAPG Datapages/Search and
1313
Discovery, article #90291.
1314
Proskurowski, G., Lilley, M.D., Kelley, D.S., 2006. Low temperature volatile
1315
production at the Lost City Hydrothermal Field, evidence from a hydrogen stable
1316
isotope geothermometer. Chemical Geology, 229 (4), 331−343.
1317
Reed, M.H., 1997. Hydrothermal alteration and its relationship to ore fluid composition
1318
In: Barnes, H.L. (Ed.), Geochemistry of Hydrothermal Ore Deposits, 3rd edition.
1319
John Wiley and Sons, pp. 303–365.
1320 1321
Rodrigues, R., 2005. Chemostratigraphy. In: Koutsoukos, E.A.M. (Ed.), Applied Stratigraphy. Springer, Dordrecht, Netherlands, pp. 165–178.
1322
Rodriguez, C.R., Jackson, C.A.L., Rotevatn, A., Bell, R.E., Francis, M., 2018. Dual
1323
tectonic-climatic controls on salt giant deposition in the Santos Basin, offshore Brazil.
1324
Geosphere 14, 215–242.
1325 1326
Roedder, E., 1984. Fluid inclusions. Mineralogical Society of America, Reviews in Mineralogy 12, 646 p.
1327
Roedder, E., Bodnar, R.J., 1997. Fluid inclusion studies of hydrothermal ore deposits.
1328
In: Barnes, H.L. (Ed.), Geochemistry of Hydrothermal Ore Deposits. Wiley, New
1329
York, pp. 657–697.
1330
Sabato Ceraldi, T., Green, D., 2016. Evolution of the South Atlantic lacustrine deposits
1331
in response to Early Cretaceous rifting, subsidence and lake hydrology. In: Sabato
1332
Ceraldi, R.A., Hodgkinson, T., Backe, G. (Eds.), Petroleum Geoscience of the West
1333
Africa Margin. Geological Society, London, Special Publication 438, pp. 77–98.
1334
Saller, A., Rushton, S., Buambua, L., Inman, K., Mcneil, R., Dickson, J.A.D., 2016.
1335
Pre-Salt stratigraphy and depositional systems in the Kwanza Basin, offshore Angola.
1336
American Association of Petroleum Geologists Bulletin 100, 1135–1164.
56 1337
Schoenherr, J., Littke, R., Urai, J.L., Kukla, P.A., Rawahi, Z., 2007. Polyphase thermal
1338
evolution in the Infra-Cambrian Ara Group (South Oman Salt Basin) as deduced by
1339
maturity of solid reservoir bitumen. Organic Geochemistry 38 (8), 1293–1318.
1340 1341
Sharp, Z., 2007. Principles of Stable Isotope Geochemistry. Pearson Prentice Hall, Upper Saddle River, New Jersey, 344 p.
1342
Sibson, R.H., 1990. Faulting and fluid flow. In: Nesbitt, B.E. (Ed.), Fluids in
1343
tectonically active regimes of the continental crust. Mineralogical Association of
1344
Canada, Short Course Handbook 18, chapter 4, pp. 93–132.
1345 1346
Simoneit, B.R.T., 1990. Petroleum generation, an easy and widespread process in hydrothermal systems: an overview. Applied Geochemistry 5 (1-2), 3–15.
1347
Simoneit, B.R.T., 2018. Hydrothermal Petroleum. In: Wilkes, H. (Ed.), Hydrocarbons,
1348
Oils and Lipids: Diversity, Origin, Chemistry and Fate. Handbook of Hydrocarbon
1349
and Lipid Microbiology. Springer, p. 800.
1350
Skinner, B.J., 1997. Hydrothermal mineral deposits: what we do and don't know. In:
1351
Barnes, H.L. (Ed.), Geochemistry of hydrothermal ore deposits, 3rd edition. New
1352
York: Wiley and Sons, pp. 1–29.
1353
Smith, L.B.J., 2006. Origin and reservoir characteristics of Upper Ordovician Trenton–
1354
Black River hydrothermal dolomite reservoirs in New York. American Association
1355
of Petroleum Geologists Bulletin 90 (11), 1691–1718.
1356 1357
Stearns, N.D., Stearns, H.T., Waring, G.A., 1935. Thermal springs in the United States. United States Geological Survey, Water Supply Paper 679-B, 59-191.
1358
Steele-MacInnis, M., Bodnar, R.J., Naden, J., 2011. Numerical model to determine the
1359
composition of H2O-NaCl-CaCl2 fluid inclusions based on microthermometric and
1360
microanalytical data. Geochimica et Cosmochimica Acta 75, 21–40.
57 1361
Suchý, V., Dobes, P., Sýkorová, I., Machovic, V., Stejskal, M., Kroufek, J., Chudoba, J.,
1362
Matejovský, L., Havelcová, M., Maysová, P., 2010. Oil-bearing inclusions in vein
1363
quartz and calcite and, bitumens in veins: Testament to multiple phases of
1364
hydrocarbon migration in the Barrandian basin (lower Palaeozoic), Czech Republic.
1365
Marine and Petroleum Geology 27 (1), 285–297.
1366
Teboul, P.A., Durlet, C., Gaucher, E.C., Virgone, A., Girard, J.P., Curie, J., Lopez, B.,
1367
Camoin, G.F., 2016. Origins of elements building travertine and tufa: New
1368
perspectives provided by isotopic and geochemical tracers. Sedimentary Geology
1369
334, 97–114.
1370
Teboul, P.A., Durlet, C., Girard, J.P., Dubois, L., Miguel, G.S., Virgone, A., Gaucher,
1371
E.C., Camoin, G., 2019. Diversity and origin of quartz cements in continental
1372
carbonates: example from the Lower Cretaceous rift deposits of the South Atlantic
1373
margin. Applied Geochemistry 100, 22–41.
1374
Teboul, P.A., Kluska, J.M., Marty, N.C.M., Debure, M., Durlet, C., Virgone, A.,
1375
Gaucher, E.C., 2017. Volcanic rock alterations of the Kwanza Basin, offshore
1376
Angola - Insights from an integrated petrological, geochemical and numerical
1377
approach. Marine and Petroleum Geology 80, 394–411.
1378
Tedeschi, L.R., 2017. Lower Cretaceous climate records and the correlation between
1379
marine and lacustrine settings (Europe and South America). (PhD Thesis). University
1380
of Oxford, Oxford, UK, p. 343.
1381
Thompson, D.L., Stilwell, J., Hall, M., 2015. Lacustrine carbonate reservoirs from Early
1382
Cretaceous rift lakes of Western Gondwana: Pre-Salt coquinas of Brazil and West
1383
Africa. Gondwana Research 28, 26–51.
58 1384
Tosca, N.J., Wright, V.P., 2014. The formation and diagenesis of Mg-clay minerals in
1385
lacustrine carbonate reservoirs. AAPG Annual Convention and Exhibition, Houston,
1386
Texas, April 6-9. USA: AAPG Datapages/Search and Discovery, article #51002.
1387
Tupinambá, M., Heilbron, M., Valeriano, C., Júnior, R.P., Dios, F.B., Machado, N.,
1388
Silva, L.G.D.E., Almeida, J.C.H., 2012. Juvenile contribution of the Neoproterozoic
1389
Rio Negro Magmatic Arc (Ribeira Belt, Brazil): Implications for Western Gondwana
1390
amalgamation. Gondwana Research 21 (2–3), 422–438.
1391
Unternehr, P., Péron-Pinvidic, G., Manatschal, G., Sutra, E., 2010. Hyper-extended
1392
crust in the South Atlantic: in search of a model. Petroleum Geoscience 16, 207–215.
1393
Utada, M, 1980. Hydrothermal alterations related to igneous activity in Cretaceous and
1394
Neogene formations of Japan. Mining Geology, Special Issue 8, 67–83.
1395
Ventura, G.T., Simoneit, B.R.T., Nelson, R.K., Reddy, C.M., 2012. The composition,
1396
origin and fate of complex mixtures in the maltene fractions of hydrothermal
1397
petroleum assessed by comprehensive two-dimensional gas chromatography.
1398
Organic Geochemistry 45, 48–65.
1399
Vieira de Luca, P.H., Matias, H., Carballo, J., Sineva, D., Pimentel, G.A., Tritlla, J.,
1400
Esteban, M., Loma, R., Alonso, J.L.A., Jiménez, R.P., Pontet, M., Martinez, P.B.,
1401
Vega, V., 2017. Breaking barriers and paradigms in presalt exploration: The Pão de
1402
Açúcar discovery (offshore Brazil). In: Merrill, R.K., Sternbach C.A. (Eds.), Giant
1403
fields of the decade 2000–2010. American Association of Petroleum Geologists,
1404
USA, Memoir 113, pp. 177–194.
1405 1406 1407 1408
White, D.E., 1957. Thermal waters of volcanic origin. Geological Society of America Bulletin 68, 1637–1658. Whitmarsh, R.B., Manatschal, G., Minshull, T.A., 2001. Evolution of magma-poor continental margins from rifting to sea-floor spreading. Nature 413, 150–153.
59 1409
Wierzbicki, R., Dravis, J.J., Al-Aasm, I., Harland, N., 2006. Burial dolomitization and
1410
dissolution of Upper Jurassic Abenaki platform carbonates, Deep Panuke reservoir,
1411
Nova Scotia, Canada. American Association of Petroleum Geologists Bulletin 90,
1412
1843–1861.
1413 1414
Wilkinson, J.J., 2001. Fluid inclusions in hydrothermal ore deposits. Lithos 55, 229−272.
1415
Wilkinson, J.J., 2010. A Review of Fluid Inclusion Constraints on Mineralization in the
1416
Irish Ore Field and Implications for the Genesis of Sediment-Hosted Zn-Pb Deposits.
1417
Economic Geology 105, 417–442.
1418
Wilson, M.I.J., Evans, M.J., Oxtoby, N.H., Nas, D.S., Donnelly, T., Thirwall, M., 2007.
1419
Reservoir quality, textural evolution, and origin of fault-associated dolomites.
1420
American Association of Petroleum Geologists Bulletin 91, 1247–1272.
1421
Wilson, N.S.F., Zentilli, M., 1999. The role of organic matter in the genesis of the El
1422
Soldado volcanic-hosted manto-type Cu deposit, Chile. Economic Geology 94 (7),
1423
1115–1136.
1424 1425 1426 1427
Winter, W.R., Jahnert, R.J., França, A.B., 2007. Bacia de Campos. Boletim de Geociências da Petrobras 15, 511–529. Wright, V.P., 2011. Reservoir architectures in non-marine carbonates. AAPG Annual Convention and Exhibition, Proceedings, Houston, p. 75–89.
1428
Wright, V.P., 2012. Lacustrine carbonates in rift settings: the interaction of volcanic and
1429
microbial processes on carbonate deposition. In: Garland, J., Neilson, J.E., Laubach,
1430
S., Whidden, K.J. (Eds.), Advances in carbonate exploration and reservoir analysis.
1431
Geological Society of London, Special Publications 370, pp. 39–47.
60 1432
Wright, V.P., Barnett, A.J., 2014. Cyclicity and Carbonate-Silicate Gel Interactions in
1433
Cretaceous Alkaline Lakes. AAPG Annual Convention and Exhibition, Houston,
1434
Texas. Search and Discovery Article #51011.
1435
Wright, V.P., Barnett, A.J., 2015. An abiotic model for the development of textures in
1436
some South Atlantic Early Cretaceous lacustrine carbonates. In: Grotzinger, J.P.,
1437
James, N. (Eds.), Microbial carbonates in space and time: implications for global
1438
exploration and production. Geological Society of London, Special Publications 418,
1439
pp. 209–219.
1440
Wright, V.P., Tosca, N.J., 2016. Geochemical Model for the Formation of the Pre-Salt
1441
Reservoirs, Santos Basin, Brazil: Implications for Understanding Reservoir
1442
Distribution. American Association of Petroleum Geologists Annual Convention and
1443
Exhibition. Calgary, Alberta, Canada.
1444
Xu, K., Yu, B., Gong, H., Ruan, Z., Pan, Y., Ren, Y., 2015. Carbonate reservoirs
1445
modified by magmatic intrusions in the Bachu area, Tarim Basin, NW China.
1446
Geoscience Frontiers 6, 779–790.
1447
Yamamoto, K., Ishibashi, M., Takayanagi, H., Asahara, Y., Sato, T., Nishi, H., Iryu, Y.,
1448
2013. Early Aptian paleoenvironmental evolution of the Bab Basin at the southern
1449
Neo-Tethys margin: Response to global carbon-cycle perturbations across Ocean
1450
Anoxic Event 1a. Geochemistry, Geophysics, Geosystems 14 (4), 1104–1130.
1451
You, D., Han, J., Hu, W., Qian, Y., Chen, Q., Xi, B., Ma, H., 2018. Characteristics and
1452
formation mechanisms of silicified carbonate reservoirs in well SN4 of the Tarim
1453
Basin. Energy Exploration & Exploitation 36 (4), 820–849.
1454
Zalán, P.V., Severino, M.C.G., Rigoti, C.A., Magnavita, L.P., Bach, J.A., 2011. An
1455
Entirely New 3D-View of the Crustal and Mantle Structure of a South Atlantic
61 1456
Passive Margin - Santos, Campos and Espírito Santo Basins, Brazil. AAPG Annual
1457
Convention and Exhibition, Proceedings, Houston, 12p.
1458
Zhu, D., Meng, Q., Jin, Z., Liu, Q., Hu, W., 2015. Formation mechanism of deep
1459
Cambrian dolomite reservoirs in the Tarim basin, northwestern China. Marine and
1460
Petroleum Geology 59, 232–144.
1461 1462
62 1463
Figures and Tables
1464 1465
Figure 1. Location map of the study area in the northern Campos Basin, offshore
1466
southeast Brazil (modified from Dias et al., 1988; Lima and De Ros, 2019). The straight
1467
dashed lines define the limits of the Campos Basin with the Espírito Santo Basin to the
1468
North (Vitória High) and with the Santos Basin to the South (Cabo Frio High). The
1469
detail indicates the location of the four Petrobras wells used in the study.
1470 1471
Figure 2. Summarized Stratigraphic chart of the Lagoa Feia Group in Campos Basin,
1472
offshore Brazil (modified from Winter et al., 2007; Lima and De Ros, 2019). Wavy
1473
lines represent unconformities.
1474 1475
Figure 3. Detailed stratigraphic logs of the Coqueiros (rift section) and Macabu (sag
1476
phase) formations at Campos Basin, offshore Brazil, in three of the four wells used in
1477
the study. The stratigraphic datum is the top of the sag reservoir (base of evaporites
1478
from Ariri Formation, which is not shown). Dolomitization and silicification intensities
1479
are indicated by the size of the bars increasing to the left. GR: gamma ray curve; δ13C
1480
and δ18O: samples analyzed for oxygen and carbon isotopes; 87Sr/86Sr: samples analyzed
1481
for strontium isotopes; FI: samples analyzed for fluid inclusions. Red wavy lines
1482
represent unconformities.
1483 1484
Figure 4. Photomicrographs highlighting main features in rift and sag lithologic types,
1485
Campos Basin, Brazil. A) Rudstone of fragmented, partially dissolved bivalve bioclasts,
1486
with well-preserved interparticle porosity (blue) (plane-polarized light with uncrossed
1487
polarizers; //P). B) Stevensitic arenite with stevensite ooids and peloids strongly
63 1488
dissolved and silicified (white) (//P). C) Fascicular calcite crusts (CFC) intercalated
1489
with granular deposits replaced by microcrystalline dolomite (MD) (crossed polarizers;
1490
XP). D) Calcite spherulites (CS) in Mg-clay matrix replaced by microcrystalline
1491
dolomite (MD) (XP). E) Planar to wavy laminite with millimetric levels defined by
1492
predominance of clay or microcrystalline dolomite (//P). F) Intraclastic grainstone with
1493
interparticle porosity partially cemented by calcite and dolomite (//P).
1494 1495
Figure 5. Paragenetic sequences interpreted for the sag fascicular calcite crusts from
1496
northern Campos Basin, Brazil. The paragenetic sequence is divided into syngenetic
1497
(Syn), eodiagenetic, mesodiagenetic and hydrothermal phases (modified from Lima and
1498
De Ros, 2019). The thickness of the lines corresponds to the intensity or abundance of
1499
the processes and products. The red lines represent porosity loss, and the blue lines
1500
porosity increase. Tha and Tho show the mean homogenization temperatures of primary
1501
aqueous and oil fluid inclusions, respectively (minimum and maximum limits in
1502
brackets).
1503 1504
Figure 6. Schematic representation of the main syngenetic, diagenetic and hydrothermal
1505
processes of formation and alteration of the studied Aptian Pre-Salt deposits, Campos
1506
Basin, offshore Brazil (modified from Lima and De Ros, 2019). (A) Syngenetic crust of
1507
coalesced fascicular calcite aggregates covering syngenetic Mg-clay matrix partially
1508
replaced and displaced by calcite spherulites. B) Microcrystalline and blocky dolomite
1509
partially replacing the Mg-clay matrix, calcite spherulites, and fascicular aggregates.
1510
Dolomite rhombohedra partially filling inter-aggregate pores. Dissolution of the
1511
syngenetic Mg-clay matrix and part of the dolomitized calcite aggregates. C)
1512
Microcrystalline quartz and chalcedony spherulites partially to fully replacing
64 1513
preexisting constituents, and filling primary and secondary porosity. D) Silicification of
1514
some dolomitized fascicular and spherulitic aggregates. Svanbergite as pseudocubic
1515
crystals disseminated in microcrystalline silica or within chalcedony spherulites.
1516
Recrystallization of fascicular calcite and initiation of fracturing process. E) Intense
1517
silicification, fracturing and dissolution. Saddle dolomite partially filling fracture and
1518
vugular pores. F) Macrocrystalline calcite, radiated Sr-barite and celestine, fibro-
1519
radiated zeolite, euhedral to subhedral pyrite, chalcopyrite, galena and sphalerite, and
1520
massive bitumen filling the secondary pores. The numbering is in accordance with the
1521
paragenetic sequence in Figure 5.
1522 1523
Figure 7. Photomicrographs showing aspects of sag lithologic types affected by
1524
dolomitization and silicification. A) Incipiently-dissolved saddle dolomite (SD) filling
1525
vugular pore (XP). B) Macrocrystalline calcite (MC) and prismatic, elongated and
1526
oriented celestine (Clt) filling fractures in brecciated dolostone (Dol) (XP). C)
1527
Cathodoluminescence (CL) image of zoned saddle dolomite crystals (SD) partially
1528
dissolved and replaced by microcrystalline silica (MS). D) Zoned saddle dolomite
1529
intensely silicified (SSD), macrocrystalline quartz (MQ), and spherulitic chalcedony
1530
(ChS) completely filling vugular pore (XP). E) Drusiform quartz (DQ) and spherulitic
1531
chalcedony (ChS) cementing vugular porosity in partially dolomitized and silicified
1532
protolith (XP). F) Radial prismatic quartz (PQ), prismatic Sr-barite (SB),
1533
macrocrystalline calcite (MC) and euhedral pyrite (blue arrows) filling vugular porosity
1534
(XP).
1535 1536
Figure 8. Photomicrographs from petrographic features of some hydrothermal mineral
1537
phases identified in this study. A) Prismatic Sr-barite (SB), saddle dolomite (SD) and
65 1538
macrocrystalline calcite (MC) (stained red) filling vugular pore (XP). B) CL image of
1539
zoned macrocrystalline calcite (MC) and saddle dolomite (SD) filling vugular porosity.
1540
C) Elongated prismatic Sr-barite (SB) and euhedral macrocrystalline calcite (MC)
1541
partially filling vugular porosity (backscattered electrons image; BSE). D) Sr-barite
1542
(SB) filling vugular porosity in dolostone (Dol) (BSE). E) Fibro-radiated zeolite (Ze)
1543
and euhedral macrocrystalline calcite (MC) filling partially vugular pore in chert
1544
composed by microcrystalline silica (MS) (BSE). F) Blocky dolomite (BD),
1545
macrocrystalline calcite (MC) and bitumen (Bi) cementing vugular porosity in rock with
1546
totally recrystallized and partially dolomitized calcite spherulites (CS) (//P).
1547 1548
Figure 9. Photomicrographs showing examples of fluid inclusions (FIs) in double-
1549
polished sections (120 µm) showing primary aqueous and petroleum fluid inclusions
1550
(blue and red arrows, respectively) from Campos Basin Pre-Salt, Brazil. Liquid-rich
1551
aqueous inclusions hosted in: A) Macrocrystalline calcite, B) Quartz, and C) Sr-barite
1552
(//P). D) Photomicrograph shows a saddle dolomite growth zones bearing aligned
1553
primary petroleum inclusions (//P). E) Photomicrograph in (//P) and F) ultraviolet
1554
fluorescence image of saddle dolomite hosting primary aqueous and pseudo-secondary
1555
petroleum fluid inclusions.
1556 1557
Figure 10. Homogenization temperature (°C) histograms of aqueous and petroleum
1558
(hatched) fluid inclusions hosted in syngenetic and diagenetic constituents from Pre-Salt
1559
sag and rift intervals, Campos Basin, offshore Brazil. (A) Primary fluid inclusions
1560
hosted in all analyzed constituents. (B) Primary fluid inclusions hosted in recrystallized
1561
calcite bioclasts. (C) Primary fluid inclusions hosted in recrystallized fascicular calcite.
1562
(D) Primary fluid inclusions hosted in recrystallized calcite spherulites. (E) Secondary
66 1563
fluid inclusions hosted in recrystallized calcite spherulites. The average formation
1564
temperature currently measured in the depth interval where the samples for fluid
1565
inclusions were obtained is 113.7 °C.
1566 1567
Figure 11. Salinities estimated for minerals phases from Pre-Salt rift and sag intervals,
1568
Campos Basin, Brazil. Histograms of the salinities (wt. % of equivalent in NaCl),
1569
computed from NaCl-H2O system, for primary and pseudo-secondary (A) and
1570
secondary (B) fluid inclusions hosted in saddle dolomite (orange), Sr-barite (black), and
1571
macrocrystalline quartz (brown) and calcite (red) hydrothermal phases. Note that the
1572
highest salinity values were obtained in the secondary fluid inclusions. The formation
1573
water salinity sampled in the rift reservoir is 27.2 wt. % NaCl eq. (*) Salinities of
1574
aqueous fluid inclusions hosted in recrystallized fascicular calcite calculated in wt. %
1575
NaCl-CaCl2 eq.
1576 1577
Figure 12. Homogenization temperature (°C) histograms of aqueous and petroleum
1578
(hatched) fluid inclusions hosted in hydrothermal mineral phases from Pre-Salt sag and
1579
rift intervals, Campos Basin, Brazil. In the left, primary and pseudo-secondary and in
1580
the right side secondary fluid inclusions obtained in all hydrothermal phases (A and B),
1581
Sr-barite (black) (C and D), macrocrystalline quartz (brown) (E and F), macrocrystalline
1582
calcite (red) (G and H), and saddle dolomite (orange) (I). Note that the homogenization
1583
temperatures obtained from oil fluid inclusions are consistently lower than those of the
1584
aqueous fluid inclusions in all mineral phases.
1585 1586
Figure 13. Carbonate isotope composition, Campos Basin Pre-Salt, offshore Brazil.
1587
Cross-plots of: A) and B) δ18O values versus δ13C values (both in ‰ VPDB) by
67 1588
lithofacies and wells. Note the gray polygons delimiting the rift (r) and sag (s) data from
1589
Campos (Dias, 1998; Rodrigues, 2005; Muniz and Bosence, 2015), Santos (Farias et al.,
1590
2019), and Kwanza Basins (Saller et al., 2016; Sabato Ceraldi and Green, 2016). C) and
1591
D) 87Sr/86Sr ratio versus δ18O values (in ‰ VPDB). E) 87Sr/86Sr ratio versus δ13C values
1592
(in ‰ VPDB). F)
1593
rocks compared to Early Cretaceous and Phanerozoic seawater (Jenkyns et al., 1995;
1594
McArthur et al., 2012; Yamamoto et al., 2013; Bodin et al., 2015; Ando, 2015) and Pre-
1595
Salt carbonate (Dias, 1998; Tedeschi, 2017; Pietzsch et al., 2018; Farias et al., 2019).
1596
Samples of microcrystalline and blocky calcite in recrystallized bioclasts (dark green
1597
triangles), recrystallized fascicular calcite (purple squares), microcrystalline and blocky
1598
calcite in laminite (light green diamonds), recrystallized calcite spherulites (blue
1599
squares), microcrystalline and blocky dolomite (pink circles), and hydrothermal
1600
carbonates filling secondary porosity and fractures (saddle dolomite - orange circles;
1601
macrocrystalline calcite - red rhombs). Micromill and bulk samples are indicated with
1602
and without external lines, respectively. Note that the δ18O values are closer to zero for
1603
primary, syngenetic and eodiagenetic constituents, more negative for the hydrothermal
1604
phases, and intermediate for the recrystallized fascicular calcite. In the 87Sr/86Sr vs. δ18O
1605
and
1606
syngenetic and eo/mesodiagenetic constituents (blue line) and another for the
1607
hydrothermal phases (red line). In the hydrothermal paragenesis, the more radiogenic
1608
values (higher
1609
well as the correlation factors (R) are indicated.
87
Sr/86Sr values obtained in the hydrothermal phases and sag host
87
Sr/86Sr vs. δ13C cross-plots there are two well-defined trends, one for the
87
Sr/86Sr ratio) present more negative values of δ18O. The equations as
1610 1611
Figure 14. Homogenization temperature (°C) versus salinity (wt. % NaCl eq.) cross plot
1612
for aqueous fluid inclusions from Pre-Salt sag and rift samples, Campos Basin, offshore
68 1613
Brazil. Primary/pseudo-secondary (upper part) and secondary (lower part) fluid
1614
inclusions hosted in hydrothermal phases Sr-barite (black), macrocrystalline calcite
1615
(red) and quartz (brown), and saddle dolomite (orange). Salinity in primary/pseudo-
1616
secondary fluid inclusions varies in the range of 12 to 26 wt. % eq. NaCl. However,
1617
97% of the data obtained from aqueous fluid inclusions show salinities higher than 15
1618
wt. % eq. NaCl. (*) Salinity of aqueous fluid inclusions hosted in recrystallized
1619
fascicular calcite calculated in wt. % NaCl-CaCl2 eq.
1620 1621
Figure 15. Burial-thermal history diagram for northern Campos Basin Pre-Salt
1622
lacustrine reservoirs, highlighting the burial depths (black line) of the sag reservoir (S)
1623
and of the 100 °C isotherm (dashed red line). The surface temperatures estimated for rift
1624
and sag deposits are 26.9 and 27.8 °C, respectively. Color scale and red values within
1625
the white rectangles represent the temperature variation. The blue arrows in the lower
1626
right corner indicate the depths of the samples analyzed for fluid inclusions. In the left,
1627
a histogram of homogenization temperatures (Th) of primary/pseudo-secondary aqueous
1628
fluid inclusions hosted in fascicular calcite, saddle dolomite, macrocrystalline calcite,
1629
mega-quartz, and Sr-barite, indicating precipitation temperatures significantly higher
1630
than the burial temperatures. Note that subsidence was essentially continuous since the
1631
deposition of the Pre-Salt sag sequence. Magmatic activity (M) occurring in the Campos
1632
Basin during Cretaceous and Paleogene (Winter et al., 2007) is shown by purple
1633
rectangles and black dashed lines. The dashed green line indicates the peak of
1634
conventional oil generation and migration (OGM) in the Miocene of the Campos Basin,
1635
according to Mello et al. (1994).
1636
69 1637
Figure 16. Schematic representation of the deep-burial hydrothermal system affecting
1638
the northern Campos Basin Pre-Salt reservoirs, offshore Brazil. The fault-focused
1639
hydrothermal system probably involved mixing of fluids derived from several sources:
1640
(B) Pre-Cambrian basement (mainly granitic-gneissic felsic rocks), (C) volcanic rocks
1641
from the Cabiúnas Formation (mainly basaltic), (R) rift deposits from the Atafona and
1642
Coqueiros formations, and (S) serpentinization of the upper mantle. Note the intrusive
1643
magmatic (mafic) rocks (M) in the Coqueiros (rift interval) and Macabu (sag section)
1644
formations; hydrothermal alteration (H) comprising extensive dolomitization,
1645
silicification, and dissolution, with the paragenesis including saddle dolomite,
1646
macrocrystalline calcite, mega-quartz, Sr-barite, celestine, fluorite, dickite, sphalerite,
1647
galena, and other metallic sulfides filling fractures and dissolution porosity; and
1648
anhydrite thick layer (A) at the top of the sag reservoir affected by the hydrothermal
1649
alteration (base of evaporites).
1650 1651
Table 1. Statistical summary of the results obtained in aqueous and petroleum
1652
inclusions from Campos Basin Pre-Salt, offshore Brazil. Fluid inclusions are hosted in
1653
the recrystallized bioclasts and syngenetic fascicular calcite, diagenetic recrystallized
1654
calcite spherulites, saddle dolomite, macrocrystalline calcite, mega-quartz, and radiated
1655
Sr-barite. Lithologic types - br: bioclastic rudstone; fcc: fascicular calcite crust; sc:
1656
stevensitic claystones with calcite spherulites; ig: intraclastic grainstone; ct: chert.
1657
Comp.: composition; FC: fluorescence color with Nikon UV-2A filter; API hc:
1658
estimated API gravity of petroleum inclusions; Th: liquid-vapor homogenization
1659
temperature of aqueous and petroleum fluid inclusions; Tm ICE: ice melting temperature
1660
of aqueous fluid inclusions; Sal (wt% NaCl eq.): equivalent salinities estimated from
1661
final ice melting temperatures in weight percent of NaCl; n: number of inclusions
70 1662
measured; pr: primary fluid inclusions; sec: secondary fluid inclusions; psec: pseudo-
1663
secondary fluid inclusions; aq: aqueous fluid inclusions; hy: hydrocarbon fluid
1664
inclusions; wt: white fluorescence; bl: blue fluorescence; yl: yellow fluorescence; pyl:
1665
pale yellow fluorescence; mx: mixed inclusions containing aqueous fluid and
1666
petroleum; N/A: could not be determined. (*) Salinities of aqueous fluid inclusions
1667
hosted in recrystallized fascicular calcite calculated in wt. % NaCl-CaCl2 eq.
1668 1669
Table 2. δ13C and δ18O results of isotopic analyses obtained in bioclasts recrystallized
1670
to microcrystalline and blocky calcite, recrystallized fascicular calcite, laminite with
1671
microcrystalline and blocky calcite, recrystallized calcite spherulites, microcrystalline
1672
and blocky dolomite replacing pre-existing constituents, and hydrothermal saddle
1673
dolomite and macrocrystalline calcite filling dissolution porosity and fractures.
1674 87
Sr/86Sr ratio obtained from recrystallized fascicular calcite and
1675
Table 3. Results of
1676
calcite spherulites, microcrystalline and blocky dolomite, and hydrothermal saddle
1677
dolomite, macrocrystalline calcite and Sr-barite filling dissolution porosity and fractures.
1678 1679
Table 4. δ18Ow in VSMOW calculated based on Th values from fluid inclusions and
1680
measured δ18Ocarb measured in carbonates from Campos Basin Pre-Salt, Brazil. The
1681
equations of O’Neil et al. (1969) and Mathews and Katz (1977) were used to calculate
1682
δ18Ow in VPDB, which was converted to δ18Ow in VSMOW by using the conversion of
1683
O’Neil et al. (1969). Details are given in item 5.2.
1684
71 1685
Supplementary Material
1686 1687
Supplementary material 1. WDS (Wavelength-Dispersive X-Ray Spectroscopy)
1688
chemical composition of carbonates from Pre-Salt reservoir in the Campos Basin, Brazil.
1689
The box-plots show the contents of (A) Na2O, (B) Al2O3, (C) SiO2, (D) MnO, (E) FeO,
1690
(F) SrO, (G) BaO and (H) Ce2O3 in mass %, obtained from fascicular (FC; purple;
1691
n=72), spherulitic (CS; blue; n=66) and macrocrystalline (MC; red; n=140) calcite, and
1692
microcrystalline (MD; light pink; n=32), blocky (BD; dark pink; n=40) and saddle
1693
dolomite (SD; orange; n=88). Thick lines and rhombs within the boxes, and circles
1694
correspond to median, mean and outliers, respectively.
1695 1696
Supplementary material 2. Scanning electron microscopy (SEM) analyses in selected
1697
hydrothermal minerals from Campos Basin Pre-Salt, Brazil. Backscattered electrons
1698
images (BSE) and As, S, Co, Cu, Zn, Ni, Fe and Pb contents (mass%) from WDS
1699
(Wavelength-Dispersive X-Ray Spectroscopy) analyses performed on sphalerite (Sp)
1700
and euhedral pyrite (Py), and Na2O, and P2O5, SO3, CaO, TiO2, FeO, SrO and BaO
1701
(mass%) analyses performed on prismatic Sr-barite (SB). A) and B) Macrocrystalline
1702
sphalerite (Sp) engulfing and replacing microcrystalline (MD) and blocky (BD)
1703
dolomite and partially filling vugular porosity (BSE); C) and D) Euhedral pyrite (Py),
1704
Sr-barite (SB), macrocrystalline calcite (MC) and prismatic macroquartz (MQ) filling
1705
vugular and fracture porosity in chert; E) and F) Vugular porosity completely cemented
1706
by prismatic Sr-barite (SB), macrocrystalline quartz (MQ), and macrocrystalline calcite
1707
(MC).
1708
Lithologic Types
Depth (m)
Comp.
Timing
FC
API hc (o)
Th (oC) range
Th (oC) average (n)
Recrystallized Calcite B Bioclast (CB)
br
XD26.0
hy
pr
wt
45-50
Recrystallized Fascicular Calcite (FC)
fcc fcc
XB93.5 XB93.5
aq hy
pr pr
wt
45-50
wt/bl wt
40-50 45-50
pr
wt
45-50
sec
wt
45-50
80-82 87-89 93-94 95-99 101-104 108.6-108.6 92-93 83-84 91-94 95-100 100-105 105.8-105.8 110-111 63-69 76-80 82-85 85-89 95-95 112.2-112.2 63-70 71-74 77.5-77.5 81-84 86-89 91.9-91.9 97.4-97.4 101-101 105-108 106-108 109-114 117-118 119-124 125-130 112-115 120-121
81.1 (2) 88.1 (5) 93.5 (2) 96.9 (5) 102.5 (5) 108.6 (1) 92.2 (2) 83.4 (2) 92.6 (4) 97.4 (8) 102.5 (3) 105.8 (1) 110.8 (2) 66 (2) 77.7 (3) 83.5 (6) 86 (5) 95 (1) 112.2 (1) 65.8 (8) 72.1 (3) 77.5 (1) 82.4 (4) 87.1 (2) 91.9 (1) 97.4 (1) 101 (1) 106.5 (2) 107 (3) 111.5 (5) 117.5 (2) 121.5 (3) 127.5 (2) 113.5 (2) 120.5 (2)
Mineral Host
Well
B B
Recrystallized Calcite A Spherulite (CS)
Macrocrystalline Calcite (MC)
C
sc
XC46.0
sig
XB71.5
sig/fcc
XB71.5/ XB98.0
fcc fcc
XB98.5 XB98.0
fcc
XB98.5
hy
aq
pr
pr/psec
Tm ICE (oC) range
Sal (wt% NaCl eq.)
-26.8
23.9-24.3 (*)
-20.9 -16.6 to -17.0 -20.1 to -20.4 -15.4 to -19.4 -18.1 to -20.8 -14.9 to -16.6 -14.8 to -17.2 -11.1 to -13.3 -16.7 to -18.4
23.0 19.9-20.2 20.5-22.4 19.0-22.0 21.0-22.9 18.6-19.9 18.5-20.4 15.1-17.2 20.0-21.3
fcc fcc
XB98.0 XB98.5
sig fcc
XB71.5 XB98.0
fcc sig
XB98.5 XB71.5
fcc
XB98.5
ct
XB91.0
fcc ct sig ct/fcc ct
XB98.5 XB91.0 XB71.5 XB91.0/ XB98.5 XB91.0
A
sc
XC46.0
C A
sig sc
C A C
sig sc sig
C
Macrocrystalline Quartz (MQ)
Prismatic Sr-barite (SB)
C
sec
hy
pr sec
aq
pr
yl yl
N/A N/A mx mx
pr/psec
sec
wt
45-50
XB71.5 XC46.0
pyl wt
33-35 45-50
XB71.5 XC46.0 XB71.5
yl wt
32-34 45-50
hy
aq
pr
pr
pr/psec
N/A 99-103 106-106 126-126 100-100 69-73 81-81 122-124 105-105 109-109 111-111 114-115 119-119 137-139 114-114 120-123 127-127 130-135 133-133 105-106 106-109 115-122 119-121 129-134 65-69.4 70-74.4 75-75 83-83 87.3-87.3 90-94 99-99 100.9-100.9 122-127 125-130 144-144 152-152 N/A 125-130 135-135
N/A (1) 101 (3) 106 (1) 126 (1) 100 (1) 71 (2) 81 (1) 123 (2) 105 (1) 109 (1) 111 (1) 114.5 (2) 119 (1) 138 (2) 114 (1) 121.5 (2) 127 (1) 132.5 (3) 133 (1) 105.5 (3) 107.5 (3) 118.5 (5) 120 (2) 131.5 (3) 68.2 (7) 73.1 (12) 75 (1) 83 (1) 87.3 (1) 92.1 (2) 99 (1) 100.9 (1) 124.5 (3) 127.5 (2) 144 (1) 152 (1) N/A (1) 127.5 (3) 135 (1)
-20.7 -15.2 to -16.1 -15.9 -16.0
22.9 18.8-19.5 19.4 19.5
-9.5 -9.0 -17.6 -15.4 to -15.9 -18.6 -17.9 to -18.0 -19.0 -19.9 to -21.2 -16.2 -16.2 to -17.8 -22.1 -16.6 to -18.2 -19.1 to -21.9 -19.5 to -21.3 -21.7 to -22.3 -19.3 to -20.5
13.4 12.9 20.7 19.0-19.4 21.4 20.9-21.0 21.7 22.3-23.1 19.6 19.6-20.8 23.7 19.9-21.1 21.8-23.6 22.0-23.2 23.5-23.9 21.9-22.7
-12.8 to -18.2 N/A -15.3 -16.6 -19.1 -18.1 to -19.6 -18.9
16.7-21.1 N/A 18.9 19.9 21.8 21.0-22.1 21.6
sec C
Saddle Dolomite (SD) C
C
sig
XB71.5
hy
pr/psec sec
br fcc br fcc/br
aq
pr
br
XF39.8 XB98.0 XF39.8 XB98.0/ XF39.8 XF39.8
br
XF39.8
hy
pr pr/psec
yl yl
yl yl
32-34 mx
N/A mx
N/A 122-124 83-83 100-100 105-107 119-119 N/A N/A 105-105 121-124 126-131 132-137 137-137 140-143 144-148 150-150 85-85 79-79 89-89
N/A (1) 123 (2) 83 (1) 100 (1) 106 (2) 119 (1) N/A (2) N/A (7) 105 (1) 122.5 (3) 128.5 (4) 134.5 (6) 137 (1) 141.5 (2) 146 (2) 150 (1) 85 (1) 79 (1) 89 (1)
-18.2 -16.9 to -19.2
21.1 20.2-21.8
-16.6 to -17.3 -19.3 to -25.3 -16.1 -18.4 to -21.0 -13.5 to -20.7 -15.0 to -17.8 -16.7 -14.5 to -16.8 -16.9 to -17.4 -20.4
19.9-20.5 21.9-26.1 19.5 21.3-23.0 17.3-22.9 18.6-20.8 20.0 18.2-20.1 20.2-20.5 22.7
Mineral Phase
Sampling
Well Depth (m) δ13C
Microcrystalline/Blocky Calcite (MBCB) (Bioclastic Grainstone and Rudstone)
bulk
B
Microcrystalline/Blocky Calcite (MBCL) (Laminite)
bulk
B
Fascicular Calcite (FC)
bulk
C B
C
micromill
C
XD06.0 XD08.0 XD10.0 XD73.0 XE38.3 XC95.4 XD00.0 XD01.0 XD43.0 XD44.0 XA15.0 XA37.0 XC85.0 XC88.2 XC89.0 XC93.0 XB57.0 XA06.0 XA95.0 XC02.3 XC04.0 XB98.8 XC01.0 XC08.5 XB83.0 XB87.0 XB88.5
XB98.8 XC08.5 Calcite Spherulite (CS)
bulk
A
XA66.5 XA84.0
1.10 0.99 1.13 0.83 0.45 1.14 0.70 1.42 1.37 1.43 1.41 0.93 1.30 1.28 1.52 1.42 1.07 1.98 1.43 1.60 1.62 1.42 1.17 1.26 0.05 1.15 0.88 0.71 0.34 0.19 0.16 0.59 0.72 0.63 0.98 0.52
δ13C error (1s) 0.04 0.06 0.05 0.06 0.05 0.03 0.05 0.04 0.04 0.05 0.05 0.05 0.07 0.05 0.05 0.02 0.05 0.02 0.09 0.05 0.04 0.05 0.05 0.05 0.06 0.04 0.03 0.04 0.05 0.03 0.06 0.04 0.02 0.04 0.05 0.05
δ18O -0.95 1.04 0.55 -1.06 -0.80 -0.61 -0.54 0.46 -1.18 -0.90 -0.59 -0.19 -0.99 -0.97 -0.23 0.63 0.54 -2.94 -1.93 -3.94 -3.34 -4.97 -4.36 -2.34 -5.70 -5.25 -5.10 -4.95 -5.37 -5.24 -5.12 -5.19 -3.85 -2.88 1.23 1.28
δ18O error (1s) 0.06 0.05 0.08 0.10 0.09 0.06 0.08 0.05 0.04 0.08 0.11 0.08 0.08 0.06 0.06 0.04 0.07 0.08 0.07 0.08 0.08 0.07 0.07 0.07 0.06 0.09 0.02 0.05 0.02 0.03 0.06 0.03 0.03 0.04 0.07 0.07
Sampling Observation Bulk analyses in the bioclastic grainstone composed of blocky and microcrystalline calcite in recrystallized bioclasts (>90%) and microcrystalline dolomite in partially dolomitized portions (<10%).
Bulk analyses in the bioclastic rudstone composed of blocky and microcrystalline calcite in recrystallized bioclasts (100%).
Bulk analyses in the laminite lithofacies constituted by recrystallized blocky and microcrystalline calcite (>90%) partially replaced by microcrystalline dolomite (<10%).
Bulk analyses in sample composed of syngenetic fascicular calcite crusts (>90%), eodiagenetic calcite spherulite (<10%) and eo/mesodiagenetic blocky dolomite (<10%).
Punctual analyses performed in the syngenetic fascicular calcite and/or fascicular aggregates.
Bulk analyses in sample constituted by eodiagenetic calcite spherulite (>90%), syngenetic fascicular calcite (<10%) and eo/mesodiagenetic blocky dolomite (<10%).
B
C
bulk
B
micromill
C
Macrocrystalline Calcite bulk (MC) micromill
C
Microcrystalline/Blocky Dolomite (MBDD) (Dolostone)
C
XB67.0 XA03.0 XA41.2 XA58.5 XA69.8 XB59.0 XC01.0 XC91.0 XA85.0 XB53.0 XD51.0 XC87.4 XC94.0 XC08.5 XB83.0 XB84.0 XB71.5 XB83.0 XB87.0 XB88.5 XC08.5
Saddle Dolomite (SD)
bulk
A
B C
XC88.5 XC96.6 XD19.0 XD19.5 XE62.0 XF17.0 XD90.0
2.43 1.25 0.75 1.79 1.49 2.14 1.54 1.91 0.31 1.36 1.67 1.25 1.16 1.44 1.37 1.28 1.18 -0.27 0.75 0.90 0.97 0.35 0.83 0.01 0.70 0.88 -1.88 -0.57 -0.04 0.22 1.03 0.34 -0.92
0.05 0.04 0.06 0.07 0.04 0.08 0.05 0.02 0.05 0.05 0.05 0.03 0.05 0.03 0.01 0.05 0.05 0.07 0.03 0.03 0.05 0.07 0.04 0.04 0.02 0.02 0.05 0.05 0.05 0.05 0.07 0.03 0.05
1.47 -2.04 0.76 -0.10 0.07 -0.56 -1.05 1.01 -2.11 -1.54 1.74 0.35 -1.12 -1.51 -1.45 -5.90 -4.92 -9.74 -9.51 -7.85 -6.19 -8.11 -7.67 -8.22 -7.82 -7.21 -6.87 -7.39 -9.49 -7.87 -7.33 -6.97 -6.35
0.07 0.07 0.10 0.08 0.07 0.12 0.09 0.08 0.07 0.07 0.07 0.10 0.07 0.04 0.04 0.07 0.07 0.07 0.03 0.04 0.05 0.07 0.04 0.05 0.02 0.02 0.07 0.07 0.07 0.07 0.07 0.11 0.07
dolomite (<10%).
Bulk analyses in the dolostone with blocky and microcrystalline dolomite (100%). Punctual analyses performed in the eo/mesodiagenetic blocky dolomite filling secondary porosity. Bulk analyses in sample composed of hydrothermal macrocrystalline calcite (>90%) and syngenetic fascicular calcite (<10%). Punctual analyses performed in the hydrothermal macrocrystalline calcite filling vugular and fracture porosity.
Bulk analyses in the dolostone and/or chert constituted by hydrothermal saddle dolomite (>90%) and macrocrystalline calcite (<10%) filling secondary porosity.
Mineral Phase
Sampling
Well Depth (m)
87
Fascicular Calcite (FC)
micromill
C
XB83.0 XB88.5 XB98.8 XC08.5
Calcite Spherulite (CS)
bulk
A
micromill
C
XA66.5 XA84.0 XB67.0 XB89.5
XB91.0
Sr/86Sr
± 2SE
Sampling Observation
0.71317 0.71329 0.71332 0.71360
0.00002 0.00002 0.00002 0.00001
Punctual analyses performed in the syngenetic fascicular calcite and/or fascicular aggregates.
0.71317 0.71338 0.71394 0.71317 0.71367 0.71365 0.71299 0.71360 0.71315
0.00001 0.00001 0.00001 0.00001 0.00001 0.00001 0.00002 0.00001 0.00001
Bulk analyses in sample constituted by eodiagenetic calcite spherulite (>90%), syngenetic fascicular calcite (<10%) and eo/mesodiagenetic blocky dolomite (<10%). Punctual analyses performed in the eodiagenetic calcite spherulite and spherulitic aggregates .
Microcrystalline/Blocky Dolomite (MBDD)
micromill
C
XC08.5
0.71351 0.71348
0.00001 0.00002
Punctual analyses performed in the eo/mesodiagenetic blocky dolomite filling secondary porosity.
Macrocrystalline Calcite (MC)
micromill
C
XB71.5
D
XD22.3
0.71136 0.71138 0.71199 0.71225 0.71215 0.71210
0.00005 0.00001 0.00001 0.00001 0.00001 0.00001
Punctual analyses performed in the hydrothermal macrocrystalline calcite (>90%) and Sr-barite (<10%) filling vugular and fracture porosity.
XD35.7 XD38.7 Saddle Dolomite (SD)
bulk
A
XC88.5 XC96.6 XD19.0 XD19.5
0.71109 0.71100 0.71123 0.71124
0.00001 0.00001 0.00001 0.00003
Bulk analyses in the dolostone and/or chert constituted by hydrothermal saddle dolomite (>90%) and macrocrystalline calcite (<10%) filling secondary porosity.
Sr-barite (SB)
micromill
D
XD35.7 XD41.6 XD42.1
0.71226 0.71209 0.71203 0.71202 0.71204
0.00001 0.00001 0.00001 0.00001 0.00001
Punctual analyses performed in the hydrothermal Sr-barite (>90%) and macrocrystalline calcite (<10%) filling vugular and fracture porosity.
Mineral Phase
δ18Ow Min
Max
δ18Ocarb Min
Fascicular Calcite (FC)
6.5
10.4
Fascicular Calcite (FC)
5.4
Calcite Spherulite (CS)
Max
Fluid Inclusion
Th (º C) Min
Max
-5.7
-1.9
Aqueous
92.2
92.2
12.4
-5.7
-1.9
Oil
83.0
111.0
6.3
16.3
-2.1
1.7
Oil
63.0
112.2
Saddle Dolomite (SD)
2.3
5.6
-9.5
-6.4
Aqueous
105.0
150.0
Macrocrystalline Calcite (MC)
3.3
11.1
-9.7
-4.6
Aqueous
101.0
130.0
Highlights A deep-burial hydrothermal system is recognized in the Campos Basin Pre-Salt carbonates. Hydrothermal phases show low δ18O values, and high Th and salinity fluid inclusions. The geochemical and fluid inclusion signatures are similar to MVT deposits. Hydrothermal fluids are the result of mixing between different sources. Atypical oil generation and migration occurred associated with the hydrothermal action.
Declaration of interests ☒ The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper. ☐The authors declare the following financial interests/personal relationships which may be considered as potential competing interests: